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Tracking Slabs Beneath Northeastern Pacific Subduction Zones
Yu Jeffrey Gu
University of Alberta, Department of Physics, CEB 348-D, Edmonton, AB, Canada, T6G 2G7.
E-mail: [email protected]
Phone: 1 780 492 2292
Fax: 1 780 492 0714
Ahmet Okeler
University of Alberta, Department of Physics, CEB 456, Edmonton, AB, Canada, T6G 2G7.
E-mail: [email protected]
Phone: 1 780 492 4125
Fax: 1 780 492 0714
Ryan Schultz
University of Alberta, Department of Physics, CEB 456, Edmonton, AB, Canada, T6G 2G7.
E-mail: [email protected]
Phone: 1 780 492 4125
Fax: 1 780 492 0714
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Abstract
Illuminating major thermal and/or compositional variations in Earth's mantle based on reflected
seismic waves is analogous to “motion tracking” in animation cinematography. Signals analyzed
by both approaches are sensitive to strong gradients in material properties and, with proper
treatments, can be used to decipher the shape or movements of the enclosed mass. In the same
spirit, this study utilizes the amplitudes of bottom-side reflected shear waves to provide first-
order constraints on the geometry and kinematics of subducted oceanic crust and lithosphere
beneath the northwestern Pacific subduction zones. The high-resolution, depth-migrated
reflection amplitudes shows large, ~1000 km wide depressions on the 660-km seismic
discontinuity, extending from the Japan sea to eastern China. The 410-km seismic discontinuity
is locally elevated by ~15 km on the oceanside of the Japan trench, where a sharp change of
transition zone thickness infers a mantle temperature increase over XX deg C. The 410-km
seismic discontinuity is locally elevated by ~15 km east of the Wadati-Benioff zone, within
which reflection amplitude drops off significantly. We further identify a strong reflector at ~530
km depth with a reflection amplitude exceeding 5% of SS amplitude. The strength of this
anomaly increases depressed with ‘avalanching’ the lower mantle west of the Hokkaido corner.
Strong correlations between the reflectivity structure and seismic velocity suggest: (1) high-
amplitude reflections generally occurs near the edges of major seismic anomalies due to strong
shear wave focusing effect, (2) ‘gaps’ in the reflection amplitudes of the 410- and 660-km
seismic discontinuities are associated with substantial topography and major mass/heat fluxes.,
and (3). The presence this reflectors residual plume(s) in this region. UNFINISHED, will work
on last.
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1. Introduction
The convergent boundary between the Pacific, Amurian, and North American plates represents
one of the fastest destruction zones of old oceanic domains. The subduction process in this re-
gion initiated during the Cretaceous times (~65-140 Ma ago) (Northrup et al., 1995; Tonegawa et
al., 2006; Zhu et al., 2010) and continues to accommodate the differential motions between the
Pacific, Eurasia, and North American plates. The deposition of old oceanic lithosphere at the
present rate of 8-9.5 cm/yr (DeMets et al, 1990; Seno et al., 1996; Bird, 2003) not only directly
influences the surrounding mantle temperature and/or mineralogy.
The morphology and spatial extent of subducted oceanic lithosphere (for short, ‘slab’) beneath
the northwestern Pacific margin have long been investigated. Among the various data types and
approaches, seismic tomography of body waves has been the most effective in constraining
details of slab geometry and surrounding mantle conditions in this region (e.g., van der Hilst et
al., 1991, 1997; Fukao, 1992; Bijwaard et al., 1998; Fukao et al., 2001; Obayashi et al. 2006;
Huang and Zhao, 2006; Zhao and Ohtani, 2009; Li and van der Hilst, 2010). Well-defined zones
of above-average P and S wave speeds have been identified along the Wadati-Benioff zone and
within the upper mantle transition zone near Korea and eastern China (e.g., Jordan, 1977; van der
Hilst et al., 1997; Widiyantoro et al., 1997; Bijwaard et al., 1998; K´arason and van der Hilst,
2000; Fukao et al., 2001; Gorbatov et al., 2000; Gorbatov and Kennet, 2002; Lebedev and Nolet,
2003; Zhao, 2004; Obayashi et al., 2006; Huang and Zhao, 2006; Fukao et al., 2009; Zhao and
Ohtani, 2009; Li and van der Hilst, 2010). The non-geometrical shape of the high-velocity zones
have inspired discussions of slab deflection toward the horizontal, which is generally referred to
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as ‘stagnation’ (Fukao et al., 1992; Fukao et al., 2001), and possible extension into the lower
mantle (see Fukao et al., 2001, 2009 for detailed reviews). The length of the flattened part of the
slab can be as large as 800-1000 km (Huang and Zhao, 2006; Obayashi et al., 2006; Fukao et al.,
2009), at least half of which can be reproduced numerically with proper treatments of trench
migration and rollback rates (). Low-velocity structures such as arc volcanism and/or
decompressional melting of stagnant slabs (Lebedev and Nolet, 2003; Zhao, 2004; Priestley et
al., 2006; Obayashi et al., 2006; Zhao and Ohtani, 2009; An et al., 2009; Wang et al., 2009; Duan
et al., 2009; Zhao et al., 2009; Li and van der Hilst, 2010; Feng and An, 2010), or hot thermal
plume(s) (Miyashiro, 1986; Ichiki et al., 2006; Zou et al., 2008; Zhao and Ohtani, 2009; Duan et
al., 2009), further underscores the wide range of dynamical processes beneath this region. These
low- and high-velocity heterogeneities can cause strong gradients in mantle temperature and/or
composition surrounding the convergent plate boundary zones
In comparison with seismic tomography, which is highly effective in resolving ‘smooth’
variations, the amplitudes and arrival-times of body waves reflected and converted at mantle
depths are more sensitive to sharp changes in rock elastic properties (Zheng et al., 2007).
Correlations between velocity and reflectivity (Shearer and Masters, 1992; Flanagan and Shearer,
1998; Li et al., 2000; Shen et al., 2008) offer greater constraints on slab geometry and dynamics
than either approach alone. For this reason, the temperature-dependent depressions on the 660-
km seismic discontinuity by 15-60 km (Shearer and Masters, 1992; Benz and Vidale, 1992; Bina
and Helfrich? Helfrich and Bina?? Li et al., 2000; Niu et al., 2005; Tonegawa et al., 2005; Shen
et al., 2008; Tauzin et al., 200??; Lawrence and Shearer, 2006; Houser et al., 2008) have been
widely cited as evidence of stagnating and ponding slab beneath the northwestern Pacific
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collision zone. The resolutions of these seismic surveys are, however, hampered by the
restrictive source-receiver distributions of converted phases and the large averaging radii in
global analyses of secondary reflected waves also known as ‘SS precursors’. In particular, while
a pioneering study of the latter phase (Shearer and Masters, 1992) provided evidence of
stagnating slab beneath the northwestern Pacific region nearly 20 years ago, further usage of
these phases in constraining detailed slab geometry and kinematics was debated (Neele et al.,
1997; Shearer et al., 1999). Discussions of the correlations between mantle reflectivity inferred
from SS precursors and seismic velocities/mantle mineralogy near subduction zones mainly
focused on broad length scales and remained qualitative (e.g., Gu et al., 2003; Lawrence and
Shearer, 2006; Houser et al., 2008).
This study analyzes a large regional dataset of SS precursors using novel processing techniques
to improve the resolution on the seismic reflectivity structure beneath the northwestern Pacific
region (Fig. 1A). The dense regional data coverage enables pre-stack depth migration that
positions weak SS precursor amplitudes at the appropriate reflection depths and locations. By
correlating reflection amplitude variations with wave speeds, we aim to provide a self-consistent,
three-dimensional (3D) snapshot of mantle reflectivity structure and deformation near the
northwestern segment of the Pacific Ocean basin. For brevity we will hereon refer to the upper
mantle transition zone as MTZ and the 410-km, 520-km and 660-km discontinuities as the 410,
520 and 660, respectively.
2. Data and method
SS precursors are a proven means for determining the depths of mantle reflectors (e.g., Shearer
and Masters, 1992; Shearer, 1993; Gossler and Kind, 1996; Gu et al., 1998; Deuss and
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Woodhouse, 2002; Flanagan and Shearer, 1998; Gu and Dziewonski, 2002; Schmerr and
Garnero, 2007; Lawrence and Shearer, 2007; Houser et al., 200XX; Rychert and Shearer, ??).
Their strong sensitivities to the reflection depth and interfacial impedance contrast beneath mid
points (see Fig. 1A), coupled with their strong sensitivity to structures away from the source and
station locations, are ideal for mapping mantle reflectivity at both global and regional scales.
We utilize all available broadband, high-gain recordings of earthquakes that took place prior to
2008. This data set is currently managed by the IRIS Data Management Center and highlights
significant efforts from GDSN, IRIS, GEOSCOPE and several other temporary deployments.
Only data from shallow events (<75 km) with magnitude (Mw) grater than 5.0 are selected for
this undertaking. The former criterion minimizes the effect of depth phase, and the subjective
magnitude cutoff ensure that source mechanism solutions are available from the Global Centroid
Moment Tensor (GCMT) project (Dziewonski and Woodhouse, 1983) for accurate computations
of PREM (Dziewonski and Anderson, 1981) synthetic seismograms. We further restrict the
epicenter distance range to 100°-160° to minimize known waveform interferences from topside
reflection sdsS and ScS precursors ScSdScS, where d denotes a discontinuity (Schmerr and
Garnero, 2007). After applying a Butterworth band-pass filter with corner periods at 12 s and 75
s to the selected data traces, we impose a signal-to-noise ratio (SNR) criterion as the ratio
between the maximum absolute amplitude of the SS and noise. The selected signal and noise
windows are (-20 sec, 60 sec) and (-170 sec, -80 sec), respectively, relative to the predicted
arrival time of SS based on PREM (Dziewonski and Anderson, 1981). All records with SNR
lower than 3.0 are automatically rejected.
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The selected transverse-component seismograms are subsequently aligned on the first major
swing of SS phase with the aid of the corresponding synthetic seismograms. As the last step of
pre-processing, we apply time shifts by the theoretical SS and S520S times through PREM
(Dziewonski and Anderson, 1981) to account for crustal (Bassin et al., 2000) and topographical
(ETOPO5 data base) variations. Since our main focus is the upper mantle transition zone, the
approximation based on SS-S520S represents an effective compromise between the 410 and 660
and may introduce an error of 3-5 km for the depth estimation of reflectors hundreds of
kilometers away from the MTZ. Generally, these model assumptions have greater impacts on the
differential times, hence reflection depths, than on the amplitudes of SS precursors (e.g., Gu et
al., 2003).
A time-to-depth migration approach, which has been previously applied to P-to-S converted
waves (Rondenay, 2009 and references therein), is introduced to convert the precursory arrivals
of SS waves to the corresponding reflection depth and location (Gu et al., 2008; Heit et al.,
2010). The SS waveforms after the corrections for crust thickness and surface topography
correspond to equalized reflection at the Earth’s surface. Hence, each time sample preceding the
reference SS time can be mapped to a crustal/mantle depth according to the predicted travel-time
tables computed based on PREM (Dziewonski and Anderson, 1981) (Fig. 1B). The sampling
rate along the depth axis is 1 km.
Finally, to obtain a 3D reflectivity image we divide the study region into uniform, rectangular
Common Mid Point (CMP) gathers with horizontal and vertical step sizes of 2° and 8°,
respectively (IS THIS TRUE, AHMET?)_. Time-to-depth migration (Zheng et al., 2007) is
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subsequently performed at each cell and the entire set of resulting migrated traces is interpolated
using a 3D, bi-linear interpolation method provided by MATLAB. Despite linear interpolation
used in each direction, the bi-linear approach constructs new data points from a discrete set of
original data values based on a quadratic function (Press et al., 1993). The resolution of this
approach is further examined in the sections below.
3. Results
3.1. Maps of Reflection Amplitudes
Fig. 2 shows the region of interest in this study. Approximately 5000 high-quality traces are
retained after the data selection procedure detailed in the previous section. The ray theoretical
reflection points of the precursors (see Fig. 2) provide adequate resolution for the entire study
area. Furthermore, the increased data coverage in the latitude range of 35°-50° facilitates a
direct comparison of the mantle reflectivity structures in the vicinity of southern/central Japan
(cross-sections A and B) with those beneath the Kuril trench (cross-section C).
The Amplitude variations of 3D depth-converted SdS waves indicate the presence of large-scale
structures in the MTZ and shallow lower mantle. The top of the MTZ (Fig. 3) contains an
elongated, highly reflective zone (HRZ), extending from the northern Great Khingan Range in
the east to the northwestern corner of the study region beneath the Gobi desert. This 1500-km
wide anomaly reaches its maximum amplitude (9% of that of SS, for short, 9%) at ~425-km
depth, which is approximately 15 km below the global average of the 410-km seismic
discontinuity (Fig. 3A) (Gu et al., 2003; Houser et al., 2008). A second, weaker HRZs is visible
east of the Wadati-Benioff zone along the Kurile and Japan arcs, peaking at ~8% amplitude near
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the Hokkaido corner (see Fig. 3A).
The HRZs at the top of MTZ decays quickly with depth and the reflectivity pattern at ~520 km
depth is dominated by a strong (5-8%), uniquely shaped reflector (Fig. 3B). The center of this
reflector is located near Sikhote-Alin Mountains, roughly coinciding with the slab corner
between Japan and Kuril subduction zones outlined by Sam Gudmundsson and Sambridge
(1998) west of the Hokkaido corner (see depth map at 540 km, Fig. 3B). The orientation of this
boomerang-shaped structure (see map at 520 km) changes from ~30 deg oblique to the trench-
perpendicular direction west of Honshu Island to trench-perpendicular beneath northeastern
China. The vertical dimension of this mid-MTZ HRZ is no greater than 40 km (see Fig. 3B).
Large-scale reflective structures are clearly visible at the base of the upper mantle (Fig. 3C) and
below (Fig. 3d). Major north-south oriented HRZs are observed at 675-km depth northwest of
the Japan-Kuril arc-arc interaction region and the eastern section of the Gobi desert, respectively
(see Fig. 3C). The maximum amplitudes of both anomalies exceed 10%. The depths of the
HRZs indicate local depressions of 20+ km on the 660 beneath northeastern China. The
geographical locations of these HRZs roughly overlap with those of two lower-mantle reflectors
detectable at 900-930 km depths. The stronger and slightly deeper of the two HRZs (see 6%
amplitude isosurface, Fig. 3D) lies beneath the slab corner between Japan and Kuril subduction
zones. This semi-linear reflective structure is approximately trench-perpendicular and spans the
entire Wadati-Benioff zone in this arc-arc interaction region.
3.2. Correlation between reflectivity and seismic velocity
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Detailed information on the temperature-dependent seismic velocity and impedance-driven
reflectivity structure is necessary to accurately characterize mantle structure and processes near
subduction zones. To explore wave amplitude vs. velocity relationship, we overlay reflectivity
depth cross-sections (Fig. 4; see Fig. 2 for reference) with high-resolution regional P velocities
reported by Obayashi et al. (2006). While the use of a regional S velocity model would be ideal,
key mantle heterogeneities in the study region are better resolved by the high-resolution P wave
tomography (see review by Fukao et al., 2009). Reflections within the depth ranges 120-150 km,
380-440 km and 630-700 km are consistently observed in all cross-sections despite substantial
lateral variations in depth and amplitude. The focus of this study is on the MTZ and lower
mantle where waveform complexities associated with SS sidelobes are minimal (e.g., Shearer,
1993; Gu et al., 2003).
Fig. 4A shows highly undulating MTZ boundaries between the Pacific Plate and the volcanic arc
near central Honshu Island. The 410 east of the Japan trench undergoes 15-20 km local
depression relative to the cross-sectional average depth of 415 km.?? This 500-km wide HRZ
reaches the maximum reflection amplitude of ~8% beneath central Honshu Island, approximately
overlapping with a P wave low-velocity zone centered between 380-400 km depths (Obayashi et
al., 2006; see also Zhao and Ohtani, 2009; Li and van der Hilst, 2010; Bagly et al., 2009). The
reflectivity structure changes sharply toward the Wadati-Benioff zone where the 410 reaches
local minima in both depth (~395 km) and reflection amplitude (~5%) (see Fig. 4, Profile A).
Complex reflective structures are also evident at the base of the MTZ east of the Japan trench.
The 660 shows 25+ km peak-to-peak topography and the undulations appear to negatively
correlate with those of the 410 along the trench dip. Major depressions are identified beneath
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eastern Sea of Japan (~680 km) and Gulf of Chihii (~673 km) (see Figs. 3C and 4, Profile A),
with the former showing a slight offset from the center of predicted MTZ high velocities.
The shape of the high-velocity structure becomes quasi-linear near northern Honshu Island
where a significant number of deep-focus earthquakes have been recorded (Fig. 4, Profile B).
The 410 remains depressed in the east of the Wadati-Benioff zone (see profile A). A strong HRZ
is visible at ~300 km depth in this region, approximately outlining with the top of the low-
velocity zone (also see Fig. 3A) above the MTZ. The reflection characteristics of the 410 are
generally consistent with those from profile A, but the lateral variations in amplitude and depth
are visibly diminished relative to the former profile. At the base of the MTZ, the 660 shows
extreme local topography in the vicinity of the Wadati-Benioff zone. The depth of the 660
beneath the island arcs is ~645 km, the shallowest level in the entire profile, which significantly
reduces the MTZ width (~225 km) along the trench dip (see Fig. 4, Profile B). This anomalous
topographic structure on the 660 is accompanied by a broad depression beneath the Sea of Japan
and Changbai hotspot. The 1000-km wide structure west of the Hokkaido corner overlaps with a
P wave high-velocity zone near the base of the MTZ, but its lateral dimension is considerably
greater than that inferred from the 1+% P velocity variations.
The high-velocity structure beneath the Kuril subduction zone (Profile C) is visibly more
complex than those beneath the Japan subduction system, providing convincing evidence for 1) a
fast zone along slab dip that extends down to 750+ km depths, and 2) a horizontal MTZ anomaly
west of the Sea of Okhotsk with a possible ‘necking’ beneath the Sikhote-Alin Mountains. The
reflectivity structure in Profile C accentuates the complex slab morphology and kinematics in
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this region. Apparent reflection gaps are observed on the prodominantly continuous 410 and 660
along the Wadati-Benioff zone, with the latter anomaly nearly spanning the entire Sea of
Okhotsk. The shape of the 660 phase boundary west of this low-amplitude region closely
matches the outline of the 1% high-velocity structure in the MTZ (see Fig. 4, Profile C). We also
identify a highly undulating, piece-wise continuous lower mantle reflector beneath this profile,
showing the largest amplitude (~10%) beneath the reflection gap on the 660. The presence of
this lower-mantle reflector and isolated MTZ HRZs will be discussed in detail in Section 4.
The cross-sections shown by Fig. 4 (Profiles A-C) paint markedly different pictures of MTZ
reflectivity structures between Japan and Kuril subduction zones. A north-south transect over the
deepest part of the Wadati-Benioff zones (Fig. 4, Profile D) highlight the key observations that
differentiate between these two subduction systems. South of Hokkaido corner, large-scale
high-velocity structures appear to reside within the MTZ. Despite slightly reduced amplitudes,
the MTZ phase boundaries are generally detected and laterally continuous. In particular, the 660
is generally deeper than regional averages and the largest ‘visible’ depressions is detected
between the Korea Strait and Sea of Japan. On the other hand, the Kuril subduction zone
embodies a vertically continuous high-velocity structure that extends into the shallow lower
mantle. This P velocity anomaly is supported by a strong HRZ at ~930 km depth. Furthermore,
the amplitudes of the MTZ phase boundaries in the same regions are clearly below the threshold
of detection using SS precursors. It is worth noting that 1+% P velocities appear to reach a
depth of ~280 km, which coincides with a strong, possibly deformed, shallow mantle reflector
between the two subduction zones.
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A common observation between the Japan and Kuril subduction zone is the presence of mid
MTZ reflector(s) (see Fig. 4, Profiles A-D). We identify a single HRZ with maximum
amplitudes in excess of 6% at ~525 km near or within the Benioff zones in the southern profiles.
Two isolated mid-MTZ reflectors are present under the Kuril subduction zone at approximate
depth ranges of 500-530 km and 580-600 km, respectively. The depths of these reflectors vary
considerably in each profile, whereas the amplitudes generally increase from South to North.
3.3. Hypothesis testing and nominal resolution
Several procedures are implemented to ensure the stability and accuracy of the migration method
as well as the resolution of the SS precursor data set. To investigate the effect of earthquake
source and the migration algorithm, we compute synthetic seismograms (Fuchs and Muller,
1971; Kind, 1978; Hermann and Wang, 1985) for all source-station pairs based on PREM
(Dziewonski and Anderson, 1981) and earthquake source information from GCMT (see also
Section 2). The synthetic data set is then subjected to the same filtering, binning and migration
procedures as the actual observations. Fig. 5 shows the sample output for Profile C, which
validates at least two key premises of this study. First, the two bounding MTZ reflectors are
migrated to 400 and 670 km, respectively, to at least two decimal places. These values are
consistent with those of PREM, the 1D model used in the migration procedure, which suggest
that the time-to-depth mapping of the actual data is precise in the absence of lateral variations in
velocity or phase boundary topography. Furthermore, the amplitudes and depths of the MTZ
phase boundaries are nearly constant along the profile, which imply that the collective influence
of earthquake source mechanism, station response, and phase equalization on the results of data
migration stacks is negligible along this (see Figs. 5) and other (not shown) profiles.
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Questions have surfaced in recent years regarding the accuracy of the structure/topography
inferred from SS precursors due to the mini-max nature of reflected waves and their wide Fresnel
zones at long periods (Neele et al., 1997; Chaljub and Tarantola, 1997). Shearer et al. (1999)
addressed some of the potential biases through a multi-scale resolution analysis. By inverting for
synthetic differential travel times, they showed that a topographic inversion using long-period SS
precursor observations is virtually immune to smaller-scale artifacts at a major subduction zone.
Recent high-resolution images from the investigations of subduction slabs (Schmerr and
Garnero, 2007; Heit et al., 2010), hot mantle plumes (Schmrr and Garnero, 2006; Gu et al., 2009;
Cao et al., 2011) and lithosphere (Rychert and Shearer, 2009) are further testimonies of the
strong resolvability of SS precursors on finer-than-expected structures at mantle depths. Shear
waves has been known to resolve structures with length scales beyond their ‘nominal’ resolution,
especially when waveform information is incorporated (Ji and Nataf, 1998; Mégnin and
Romanowicz, 2000). In the case of SS precursors, minor errors are expected when relatively
large Fresnel zones of SS precursors collapse onto the fine grid adopted by this study, though the
lateral depth/amplitude differences between the averaging centers could persist and the apparent
connections between reflection amplitude, seismic velocity and seismicity (see Sections 3.1 and
3.2) are hard to dismiss as random occurrences.
Without repeating the successful experiment performed by Shearer et al. (1999), we examine
different CMP sizes to determine the optimal level of tradeoff between stability and resolution.
Fig. 6 shows the a comparison of reflectivity maps at 680 km based on averaging bins sizes of 2
x 6 = 12 deg2 (Fig. 6A) and 5 x 10 = 50 deg2 (Fig. 6B). Differences in the suggested spatial
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scales of the anomalies are apparent. A significant number of reflectors, some poorly resolved
due to insufficient data, exist in the former map whereas larger bin sizes tends to over-damp the
lateral variations in 660 topography. However, the location and maximum amplitudes of major
HRZ, e.g., a semi-linear structure across northern Honshu Island and a large, uniquely shaped
zone contouring the deepest part of the arc-arc interaction region, are minimally affected by bin
sizes. Our final choice of averaging area (32 deg2) represents an effective, albeit subjective,
compromise between image stability and resolution.
3.4. Uncertainty of reflectivity structure
We estimate the uncertainty of the reflectivity profiles based on bootstrapping resampling
algorithm (Efron, 1977). For each averaging bin, we first construct a ‘bootstrapped’ data set of
equivalent size to the original data set through random drawing. This procedure is performed
with the aid of a random generator (Press et al., 1992) and allows for repeated selections of the
same seismogram. We then perform data stacking and migration on this simulated data set and
obtain a single summary migrated seismogram for this particular averaging bin. This random
drawing and migration/averaging procedure is repeated 300 times in the same data gather to
obtain a statistically significant distribution of reflectivity at each depth. We estimate the
effective uncertainty by the standard deviation of these 300 bootstrapped seismograms (Efron,
1977; see also Shearer, 1993; Deuss and Woodhouse, 2002; Gu et al., 2003; Lawrence and
Shearer, 2006; An et al., 2007; Zheng et al., 2007), and apply the same treatment to all averaging
bins along each profile.
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The bootstrapped reflectivity profiles, which are constructed based on the average of the re-
sampled seismograms at each data gather, are nearly identical to the respective profiles shown by
Fig. 4. The bootstrapped uncertainties based on one standard deviation (Fig. 7) are generally
lower than 3% below 150 km. The spatial variation in uncertainty is nearly random, which
implies that the main MTZ reflectivity structures are reasonably well resolved in all profiles.
However, all four profiles show a 200-500 km wide section of increased uncertainties (reaching
~3% amplitude) that intercepts the seismogenic zone, e.g., beneath the Japan trench in Profiles A
and B and Strait of Tartary in Profile D (see Fig. 7). This anomalous zone is partly caused by
relatively sparse data coverage (see Fig. 2), though the scattering associated with inclined high-
velocity slab structures cannot be ignored. Further discussions of the latter effect will be
provided in Section 4.
4. Interpretation and discussion
Using reflected/scattered waves to illuminate the shape of major thermal and/or compositional
anomalies is analogous to ‘motion tracking’ in animation cinematography. In a nutshell, both
procedures take advantage of the relationships between reflection/scattering strengths and
changes in material properties including density, bulk or shear modulus and, in the case of
motion tracking, index of refraction of electromagnetic waves. Signals analyzed by both
applications are strongly sensitive to gradients in material properties and, with proper treatments,
can be used to decipher the shape or movements of the enclosed mass. On the other hand,
destructive interference or scattering of the waves caused by structural asperities could present
challenges, albeit providing additional information, to both applications. The incorporation of
additional physical constraints could be highly beneficial. For the case reflectivity imaging, the
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combination of reflectivity imaging and seismic tomography can substantially improve our
existing knowledge on morphology of subducted crust and lithosphere in the northwestern
Pacific region.
Many important factors must be considered in the discussion of the morphology and kinematics
of subducting slabs. From a mineralogical viewpoint, the slab geometry and the width of the
MTZ are strongly influenced by mineralogical phase transformations of olivine to wadsleyite
(near 400 km), wadsleyite to ringwoodite (near 520 km), and ringwoodite to perovskite +
magnesiowustite (near 660 km) (Katsura and Ito, 1989; Ita and Stixrude, 1992; Helffrich, 2000;
Bina, 2003 and references therein; Akaogi et al., 2007). The endothermic phase change at the
base of the MTZ increases local buoyancy forces, which can deflect subducting slabs and aid its
stagnation within the upper mantle (Christensen, 1995; Billen, 2008, 2010; Fukao et al. 2009).
Under thermodynamic equilibrium, a cold, water-rich slab is expected to raise the 410, depress
the 660 (due to the opposite signs of their Clapeyron slopes), and be responsible for a wide range
of reflective bodies within the mantle. The presence of water can strongly impact the phase
changes in the MTZ (e.g. Inoue et al. 1995; Kohlstedt et al., 1996; van der Meijde, 2003; Ohtani
et al. 2004; Kombayashi and Omori, 2006; Huang et al., 2006; Litasov et al., 2006; Suetsugu et
al., 2006). Below is a detailed account of some of the observed reflectivity structures in the
general framework of MTZ mineralogy and temperature.
4.1. Amplitudes of the MTZ discontinuities
The amplitudes of the reflections from the MTZ phase boundaries are functions of the impedance
contrast across the reflecting surface and the transition width. Furthermore, due to the
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summation of multiple seismograms at each location and the use of SS amplitude as the
normalization term, the topography on the interface and regional variations of SS can also
significantly impact the relative SS precursor amplitudes. This study exclusively focuses on the
positive reflections associated with increased material impedances with depth. This is a
subjective decision prompted by the simple observation that the signs of well-resolved
reflectivity structures are predominantly positive in our study area. Admittedly, many positive
phases are accompanied by sizeable negative peaks that could result from reductions in velocity
and/or density, e.g., near the top of a low velocity zone or the bottom of a high velocity structure.
We defer discussions of negative phases to a future study.
The detectable ranges of amplitudes are 4-9% for S410S and 4-12% for S660S, both showing
significant lateral variations. The former range overlaps with the predicted values of ~8% from
PREM (Dziewonski and Anderson) and global average of 6.7% (Shearer, 1996) based on SS
precursor observations, whereas the latter range falls well short of the predicted 14% (Shearer,
2000). These individual amplitude estimates are strongly affected by the strength of SS, the
normalizing reference phase. For instance, the presence of attenuating low-velocity structures
(e.g. Zhao et al., 1992, 1997, 2004; Lei and Zhao, 2005; Huang and Zhao, 2006), especially near
back arc regions (e.g., Xu and Wiens, 1997; Roth et al., 1999, 2000), could reduce the absolute
amplitude of SS and increase the relative amplitude. Compositional variations associated with
Al at the base of upper mantle (e.g., Weidner and Wang, 1997, 2000; Deuss and Woodhouse,
2002; Deuss, 2009) or Fe content (Akaogi et al., 2007; Inoue et al., 2010) are also known to
broaden phase boundary widths and cause reductions in precursor amplitudes. A more stable
parameter is the amplitude ratio between the 410 and the 660 (e.g., Shearer, 2000), which we
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estimate to be within the range of 0.7-0.8. This value is slightly higher than the earlier estimates
of 0.64-0.68 based on global SS precursor (shearer, 1996) and regional ScS observations
(Revenaugh and Jordan, 1991), but it is in poor agreement with that of PREM (0.5). A regionally
sharp 410 (e.g., Benz and Vidale, 1993; Vidale et al., 1995; Neele, 1996; Melbourn and
Helmberger, 1998; Ai and Zheng, 2003; Jasbinsek et al., 2010) could , although the presence of a
fluid-rich lens near the 410 (Smyth and Frost, 2002; van der Meijde, 2003; Inoue et al., 2010).
While these effects are difficult to constrain reliably based on seismic observations, scattering
associated with undulations on the two MTZ bounding discontinuities are more readily
observable (Shearer, 2000). The presence of dipping structures, particularly in the vicinity of
slabs, can preferentially lower the ‘perceived amplitude’ of the 660, hence the amplitude ratio of
410 vs. 660, due to the 25-30% larger topography on the 660 relative to that on the 410 (see Figs.
3 and 4). The following sections carefully examine discontinuity depths and their implications
for slab geometry and dyanmics.
4.2 Depth correlation of the MTZ discontinuities
The migrated reflectivity profiles provide new insights on the effect of mantle temperatures on
phase boundary variations. Results from high-pressure mineral physics (e.g., Katsura and Ito,
1989; Ita and Stixrude, 1992; Irifune et al., 1998; Helffrich, 2000; Akaogi et al., 2007) have
predicted a negative correlation, hence an increased transition width, between the phase
boundary undulations in an olivine-dominated mantle. Seismic evidence from regional (e.g., Li
et al., 2000; Collier et al., 2001; Lebedev et al., 2002; Saita et al., 2002; Ai et al., 2003; van der
Meijde et al., 2005; Ramesh et al, 2005; Tonegawa et al. 2005) and global (Shearer and Masters,
1992; Shearer, 1993; Gossler and Kind, 1996; Gu et al., 1998; Flanagan and Shearer, 1998;
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Lawrence and Shearer, 2006; Houser et al., 2008) analyses have generally supported this
hypothesis, but analyses based on lower-resolution approaches have largely attributed the
increased thickness to a strongly deformed 660 that correlates with the thermal variations at the
base of the upper mantle (Flanagan and Shearer, 1998; Gu et al., 1998, 2003; Gu and
Dziewonski, 2002; House et al., 2008). The depth of the 410 remains problematic in view of
mantle chemistry on all scales (e.g., Gilbert et al. 2002; Fee & Dueker 2004; Du et al. 2006; Gu
and Dziewonski, 2002; Gu et al., 2003; Deuss, 2007; Schmerr and Garnero, 2007; Tauzin et al.,
2008). Additional assumptions involving corrections (Flanagan and Shearer, 1998; Gu et al.,
2003; Schmerr and Garnero, 2006; Deuss, 2007; Houser et al., 2008) and/or mechanisms
predicated on extensive compositional variations (Schmerr and Garnero, 2007; Deuss, 2007; Gu
et al., 2009; Houser and Williams, 2010) are needed to reduce the difference between observed
and expected MTZ phase boundary perturbations.
To examine the correlation between temperature and discontinuity topography in our study area,
we focus on Profile A where both the 410 and 660 show the largest detectable topographic
variations and amplitudes near the Wadati-Benioffz zone (Fig. 8). The respective peak-to-peak
depth variations of the 410 and 660 are approximately 30 km and 410 km, which are comparable
to the largest variations reported by earlier global studies (Shearer, 1993; Gossler and Kind,
1997; Flanagan and Shearer, 1998; Gu et al., 2001, 2003; Houser et al., 2008; Lawrence and
Shearer, 2008). Both phase boundaries undergo extreme deformation from the trench onset to
the deepest part of the Wadati-Benioff zone across southern Japan (see Fig. 8A). A simple bin-
by-bin correlation assuming vertical thermal structures, the same approach used in the
aforementioned global studies, suggests a positive correlation between discontinuity depths over
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the length of the profile (see Fig. 4). To account for non-vertical structures following the slab dip
(~30 deg, Gudmundsson and Sambridge, 1998), we revise the correlation analysis by applying an
indexing change such that the depth of the 410 at a given location is correlated with the 660
depth at a location ~200 km further inland. The dip-corrected phase boundaries show clear
negative correlation in the vicinity of the slab (see Fig. 8A) and the corrected correlation
coefficient is -0.4 for the entire profile, a statistically significant value that clearly favors a
thermal origin for the observed MTZ topography. A key reason for the strong negative
correlation is the observed elevation of the 410 within the Wadati-Benioff zone. This feature
represents a major departure from those of earlier time-domain global studies of SS precursors
(e.g., Flanagan and Shearer, 1998; Gu et al. 2003), which we attribute to improved data
resolution in this study. From a broader perspective, this experiment not only highlights the
ability of SS precursors in resolving small-scale subduction zone anomalies, but also provides a
blueprint for to improve global correlation analyses via a priori information such as slab dip
angles.
4.3. Continuity of the 410 beneath northeast China
There have considerable discussion of results obtained from laboratory experiments on the
existence and support for a water/melt rich layer near the top of the MTZ (Wood, 1995; Inoue et
al., 1995, 2010; Kohlstedt et al., 1995; Smyth and Frost, 2002; Frost and Dolejs, 2007). Based on
these studies, wadsleyite has a strong capacity to accommodate hydroxyl (OH−), storing up to 3
wt.% H2O under equilibrium conditions (Wood, 1995; Inoue et al., 1995, 2010; Smyth and
Dolejs, 2007). These laboratory-based measurements have been supported by regional (e.g.,
Revenaugh and Sipkin, 1994; Zheng et al., 2007; Schmerr and Garnero, 2007; Schaeffer and
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Bostock, 2010) and global (Tauzin et al., 2010) seismic observations of low-velocity zones at
similar depths that cannot be sufficiently explained by thermal variations. The infiltration of
hydrous melt is further constrained through geodynamical calculation and synthesis (e.g.,
Bercovici and Karato, 2003; Karato, 2006; Leahy and Bercovici, 2007, 2010).
Our migrated reflectivity structures provide further regional constraints on this hypothesized
hydrous layer above the 410. The 410 west of the Wadati-Benioff zone (Fig. 8B) is consistently
shallower than the regional average in this study. The largest topography is observed in the
southernmost cross-section, reaching a depth of ~400 km beneath Korea and northeastern China.
The two northern profiles B and C show modest highs of ~410 km in the topography of the 410
near the Changbai hotspot and Sikhote-Alin Mountains, respectively. The average amplitudes of
the 410 in all three profiles far exceed the regional average, despite visible falloffs in the middle
of the highlighted section in the latter two profiles (see Fig. 8B). These characteristics are
reminiscent of those reported beneath the Tonga subduction zone (Zheng et al., 2007) based on
migrations of precursors to both P and S depth phases. However our highlighted section shows
strong positive reflections, which is opposite to those reported near Tonga, and the perturbations
in depth (<15 km relative to 410 km) is weaker than those presented by the earlier study (>20
km). Our highlighted region (see Fig. 8B) is also farther away from the Wadati-Benioff zone
than the target area in Zheng et al. (2007), though metasometism involving slab-derived fluids
rising through the flattened part of slabs (see Fukao et al., 2009 for review) could potentially be
as extensive as that beneath slab wedge. In fact, intraplate volcanoes nears Changbai mountains
and Wudalianchi region (see also Fig. 8B, Profile B) have been closely linked to processes
similar to back-arc spreading of the Japan slab (Lei and Zhao, 2005; Huang and Zhao, 2006).
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Schmerr and Garnero (2007) present another intriguing comparison. Based on multiple cross-
sections in South America, this earlier study inferred a ‘melt lens’ based on evidence of delayed
and split/missing S410S reflections east of the Nasca-South America convergent zone. The
presence of highly anomalous underside reflections received further support from Contenti et al.
(submitted, 2011) based on the method presented in this study. However, the complexity of the
S410S signal from South America far exceeds that from northeastern China. Should a fluid-rich
layer be present atop the MTZ beneath our study region, its spatial scale, infiltration/storage
mechanism and/or chemistry are likely to be different from those near Tonga and South America
subduction systems.
4.4. Slab stagnation and distortion
Subducted ocean basins in the western Pacific region have been known to deflect to a near-
horizontal direction the MTZ for nearly two decades (Okino et al., 1989; van der Hilst et al.,
1991; Fukao et al., 1992, 1993). Since then, ample evidence of slab stagnation (Fukao et al.,
1993, 2001) in subduction zones worldwide has been provided by global and regional
tomographic images with improved accuracy and resolution (Fukao et al., 2001, 2009; Zhao and
Ohtani, 2009; Li and van der Hilst, 2010; Sugioka et al., 2010) and anomalous dip-angle
variations suggested by the distribution of intermediate-depth earthquakes (Chen et al., 2004).
The conditions and characteristics of stagnant lithosphere have been constrained further through
numerical calculations incorporating thermo-petrological buoyancy forces (Tetzlaff and Schmeling, 2000;
Bina et al., 2001; Bina and Kawakatsu, 2010), rheology (Billen and Hirth, 2007; Billen, 2008), and plate
history and rollback (Torii and Yoshioka, 2007; Christensen, 2010; Zhu et al., 2010).
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With the help of seismic velocities, the reflectivity information provided by our study can place
crucial constraints on slab deformation at the base of the MTZ and the shallow lower mantle. In
particular, the shape of the HRZs near the 660 provides useful measures for the geometry and
dimension of the stagnant slabs. The two southern profiles presented in Fig. 3C and Fig. 4A-C
consistently show two distinct zones of large-lateral scale depression (Fig. 9), 1) near the
piercing point of the slab at the base of upper mantle, and 2) in the second half of the stagnant
slab inferred from recent tomographic models (e.g., Huang and Zhao, 2006; Fukao et al., 2009).
The two depressive zones have nearly identical shapes, particularly in Profile B, and depth of the
660 between them ranges from 655 to 660 km in both cases. Profile A shows significantly larger
topography than Profile B near the slab piercing point. For an isochemical mantle, the maximum
depth of ~685 km would suggest a temperature increase of XX-XX deg C depending on the
selected Clapeyron slope (REF). The reduction in topography from south (Profile A) to north
(Profile B) along the island arcs is in general agreement recent studies based on receiver
functions (Niu et al., 2005) and postcursors to sScS (Yamada and Zhao, 2007). The reduced
horizontal gradient in the topography of the 660 beneath northern Honshu could be caused by a
‘soft’ slab (Li et al., 2008) under the influence of trench migration and rollback. However, Li et
al. (2008) detected little or no oceanward broadening of the 660 from high-resolution S to P
converted waves. This is inconsistent with the apparent shift between the high-velocity contours
and the onset of the depressive zones in the vicinity of the island arcs (see Fig. 4 and Fig. 10).
Resolution differences of the two data sets (SS precursors vs. receiver functions) may be a
contributing factor, still, 100-300 km horizontal broadening/ponding of the Pacific slab at the
base of MTZ in the oceanward direction remains a strong possibility.
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A dimensional analysis of slab geometry based on the topography on the 660 is informative but
requires subjective definitions. Assume the points of intersection at 670 km depth mark the
corners of the topographic structures, we estimate the horizontal dimensions of depressive zones
to be 350-450 km in Profile A and 500-600 km in Profile B. The respective topographic highs
between the depressions are estimated to be ~700 km and ~400 km. The total length beyond the
depressions near the slab piercing point is approximately 1050 km for Profiles A and 900 km for
Profile B. These values are reasonably consistent with the estimated length of 800-1000 km for
deflected slab bodies (Huang and Zhao, 2006; Fukao et al., 2010), especially if slight reductions
due to horizontal averaging are considered in our estimates. However, as suggested by Fig. 10
and the estimates above, the truly ‘flat’ part of the slab that depresses the 660 phase boundary is
most-likely less than 600 km in width.
The migration-based topography of the 660 (see Fig. 9) challenges the ‘flatness’ of stagnant
slabs. The observation of contention is the average or shallow 660 between the depressive
zones, particularly in Profile A, whereas broad, continuous depression zones have been reported
earlier though seismic tomography (see Fukao et al., 2009 for review) and reflection depth/MTZ
thickness imaging (e.g., Shearer and Masters, 1992; Flanagan and Shearer, 1998; Gu et al.,
1998, 2003; Lawrence and Shearer, 2006; Houser et al., 2008). Furthermore, the amplitude of
the 660 within this uplifted region is consistently higher than the regional averages, which is
consistent with the expected decrease of ringwoodite-perovskite+magnisiowustite phase loop
under high-than-average temperatures. The observed phase boundary behavior is plausible
based on recent geodynamical calculations of slab geometry that consider 1) trench retreat
(Christensen 1996; Tagawa et al. 2007; Zhu et al., 2010) or 2) temperature- and pressure-dependent
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viscosity (Karato and Wu 1993; see Fig. 12 of Fukao et al. 2010). These calculations infer distinct
zones of depression at the slab piercing and re-entry points, between which the 660 remains
largely unperturbed. The images provided by these models are consistent with our observations
in the MTZ, though the expected reflections from the horizontally oriented slab segment in the
shallow lower mantle (e.g., Fukao et al., 2009) are not clearly observed from our data set (see
Fig. 9).
Alternatively, the internal undulations within stagnated slab body could suggest vertical
deformation of slab interface in the MTZ. Part of the lateral variations may be related to
advection (Kellogg et al. 1999; Obayashi et al., 2006), where the ambient and relatively hot
mantle material got ‘trapped’ during the interaction between the tip of the downgoing slab and
viscous lower mantle. Trench migration and rollback history could play a major role, as the
current geometry of stagnant slab could reflect changes in slab dip over the course of 100+ Ma
(see Schmid et al., 2002 for the case of Farallon plate subduction). Finally, the presence of water
(e.g., Listov et al. 2002, 2006; Inuoe et al., 2010) and possible separation of oceanic crust from
the downgoing lithosphere (Irifune and Ringwood, 1995;van Keken et al., 1996; Hirose et al., 1999,
2005) could also contribute to strong gradients in the topography of the 660 within the ‘flat’ part of the
slab.
4.5. Slab penetration beneath Kuril subduction zone
The reflectivity structures add new insights into the long-standing debate about the depth of slab
in the Pacific northwest (van der Hilst et al., 1991; Fukao et al., 1992; van der Hilst et al., 1997;
Fukao et al., 2001, 2009). While the vertical extent of slabs and the general style of mantle
convection remain debated on the global scale, there is growing evidence of scattered and
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deformed slab material in the lower mantle (van der Hilst et al., 1997; Bijwaard et al., 1998;
Fukao et al., 2001, 2009; Obayashi et al., 2006; Courtier and Revenaugh, 2008; Li and van der
Hilst, 2010; Chang et al., 2010).
Among the various HRZs documented in this study, MTZ anomalies contained in Profiles C and
D provide strong evidence for penetrating slabs in the western Pacific region. The most visible
change in the reflectivity structures from central Honshu slab to southern Kuril slab is the
amplitude reduction of the 410 and 660, highlighted by the apparent reflection gaps in Profiles C
and D. These gaps coincide with the Wadati-Benioff zone of the Kuril slab and their lateral
dimensions reflect the increasing width of the high velocity structure from the top to the bottom
of the MTZ (see Fig. 10A). The origin(s) of these reflection gaps remain(s) debatable. Factors
that have considerable impact on the amplitudes of the MTZ reflectors (see also Section 4.1)
include Al, water and Fe contents and optics.
There are merits and significant caveats in attributing the observed reflection gaps to variations
in mantle chemistry (e.g., the first three factors listed above). Under proper mantle conditions,
an increase in Al content could broaden the depth range of garnet-to-perovskite transformation
and influence olivine and pyroxene normaltive proportions near the base of the upper mantle
(Gasparik, 1996; Weidner and Wang, 1998; 2000). In a low temperature regime, e.g.,
subduction zones examined in this study, majorite garnet (a Al bearing mineral group) can
transform to metastable ilminite that eventually transforms to Ca-perovskite (e.g., Weidner and
Wang, 1998). These phase transitions exhibit different phase boundary behaviors from the
olivine system and adversely impact the interpretation of discontinuity depths and amplitudes.
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The presence of Al-bearing Akimotoite could introduce further complexities, e.g., a high velocity
layer or a steep velocity gradient, to mid MTZ depths at low temperatures (Gasparik, 1996;
Wang et al., 2004). However, changes in Al content mainly impact mantle reflectivity structure
under mid-to-lower MTZ pressure-temperature conditions (e.g., Weidner and Wang, 2000; Wang
et al., 2004). The restrictive condition greatly weakens the role of Al in view of the unexplained
absence of the 410 within Kuril slab.
Water transported into the MTZ by the subducting slab could also modify the impedance
contrast, hence the visibility of a reflecting body (van der Meijde, 2003; Ichiki et al., 2006).
Aided by strong capacities of wadleyite and ringwoodite to retain water (Inoue et al. 1995, Kohlstedt
et al. 1996; see Fukao et al., 2009 for review), a hydrous MTZ can simultaneously affect the width
and depth of the 660 (Litasov et al., 2006; Akaogi et al., 2007; Inoue et al., 2010). However, the
effect of water on the phase phase loop of the olivine-Wadsleyite transition is rather complex and
relatively minor with1 wt% H2O (Inoue et al., 2010). The implication is that a large amount of
water must be present in the descending slab to diminish the amplitude of S410S below the
detection threshold. Unfortunately, recent seismic observations (Fukao et al., 2009; Bina and
Kawakatsu, 2010), particularly those based on a novel modeling strategy for MTZ water content
(Suetsugu et al., 2006, 2010), have largely inferred ‘dry’ (e.g., <0.5%, Suetsugu et al., 2010)
slabs in various parts of the Pacific rim. Mechanism(s) predicated on increased Fe content in
slabs are similarly flawed. While increasing the Fe number can substantially broaden the phase
loops of both olivine-wadsleyite and ringwoodite-perovskite+magnisiowustite transitions
(Litasov et al., 2006; Akaogi et al. 2007; Inoue et al., 2010), the observational support for the
enrichment of Fe in subduction zones is not well established.
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The observed reflectivity gaps are best explained by effects commonly observed in optics.
Similar to the scattering of light, the observed amplitudes of the underside SH-wave reflections
are strongly influenced by the geometry of the reflecting surface. A dipping structure or
interface generally causes defocusing or scattering that, depending on the size of the structure
relative to the wavelength of the incoming wave, can result in the destructive interference of the
reflected/scattered waves. Therefore, local topography on the two MTZ bounding phase
boundaries in response to thermal and/or compositional variations are expected to tradeoff with
reflection amplitude obtained through averaging. This effect was documented by Chaljub and
Tarantola (1997) based on results from finite-difference modeling of S660S amplitude in
response to local topography and higher-than-average velocities, though the conclusions of that
study has been a subject of considerable debate (e.g., Shearer et al., 1999). We hereby quantify
the relationship between topography and SS precursor amplitude based on simulations of stacked
SS precursors from a depressed zone assuming uniform (case 1, Fig. 10A) and more extreme
(case 2, Fig. 10A) spatial distributions of reflection points. Reflectivity synthetic seismograms
(Randall are computed for common explosive source recorded by a station at 130-deg epicentral
distance. This experiment is repeated for depth perturbations (positive for the 410 and negative
for the 660) ranging from 0 (unperturbed PREM model) to 40 km. The resulting stacked
waveforms of SS precursors show a steady decay with increasing vertical topography,
particularly for case 2 where the reflection-point distribution is sparse (Fig. 10A and 10B). For
both cases, the amplitude drops to 50% for undulations of 15-25 km on the 410 and 25-35 km on
the 660, which will be problematic during the detection of large topographic features. Between
the two phase boundaries, the influence of btopography is larger for the 410 than the 660 due to a
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smaller assumed velocity jump at the former interface (see Fig. 10B). The amplitude decay
could be more severe for sparsely populated data (see case 2 simulations, Fig. 10). Furthermore,
the presence of large topography can significantly modify the waveform characteristics of the
superimposed seismogram. The wave shape broadens within increasing topography and,
depending on the frequency, can split into separate low-amplitude arrivals reflecting the top and
bottom of the topographic structure, respectively (see Fig. 10).
An underpinning message from Fig. 10 is that the maximum depth of the 660 could be 700 km or
deeper in the Pacific northwest (e.g., Revenaugh and Jordon, 1989; Niu et al., 2005). Based on
the impedance contrasts suggested by PREM, the amplitudes of both phase boundaries could
easily fall below the detection threshold of ~4% during the migration procedure when the
topography exceeds 35 km for the 660 and 20 km for the 410. While this is the ‘worst case’
scenario that assumes the averaging bin size is equivalent to the surface area of the topographic
structure, it does provide a viable explanation for the missing 410 and 660 within the Kuril slab.
The waveform splitting phenomenon (see Fig. 10) also has significant implications for the
detection of double reflectors. For example, results from high-pressure mineral physics (e.g.,
vacher et al., 1998; Weidner and Wang, 1998, 2000; Akaogi et al., 2002) have provided solid
laboratory evidence for garnet-ilmenite-perovskite transition near the base of the upper mantle.
Within low-temperature slabs, these garnet-related transitions are expected to take place over 60-
100 km range in depth (Vacher et al., 1998; Akaogi et al., 2002) that are capable of generating
mild reflections in seismic waves. Observationally, the occurrences multiple reflectors have
been reported under different tectonic settings (e.g., Deuss and Woodhouse, 2002; Ai and Zheng,
2003; Tibi et al., 2007), but their presence beneath northwest Pacific have been questioned
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(Lebedev et al., 2002; Tonegawa et al., 2005; Niu et al., 2005). In this study, only Kuril slab
(Profiles C and D, Fig. 4) show strongly dipped, weak reflecting bodies centered at ~700- and
780-km depths along the slab dip. These minor reflectivity structures are barely detectable,
showing ~4% amplitude each. While it is tempting to link these secondary structures to multiple
phase transitions, our numerical experiment above also cautions that the waveform complexities
associate with steep topographic structures should be considered in the interpretations.
The presence of a high-amplitude lower mantle HRZ beneath Kuril slab (Fig. 11) provide
potentially crucial support for the vertical extension of Kuril slab beyond the 660. Phase
transitions of Ca-perovskite (Stixrude et al., 2007), metastable garnet (Kawakatsu and Niu, 1994;
Kubo et al., 2002), as well as transformations of dense hydrous magnesium silicates under lower-
mantle pressure-temperature conditions, have been suggested as the origins of a series lower-
mantle reflectors (Shieh et al., 1998; Ohtani, 2005; Richard et al., 2006 and references therein).
The association of lower-mantle reflectors with phase changes is partially supported by the local
maxima of reflection amplitude beneath the reflection gap on the 660. However, reflections
from a sub-horizontal lower-mantle HRZ in northeast China between 850-1000 km depths
present a potential counter argument. The existence of a chemical boundary (Wen and
Anderson, 1997), which would influence the convective flow of mantle, cannot be ruled out.
We interpret the presence of the lower mantle reflector as an integral part of an ‘avalanching’
slab (Tackley, 1993) based on the following observations: 1) slab gaps at the 410 and 660 that
imply substantial mass and heat flux, 2) correlated fast velocity structure that maintains a strong
amplitude to depths comparable to that of the lower mantle reflector, 3) the presence of a strong
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(if not the strongest) lower mantle reflector in the vicinity of the slab gap. These observations
are self-consistent and could result from the same process (i.e., slab penetration) under different
pressure-temperature conditions and, possibly, mantle chemistry. Since the lower mantle HRZ
resides directly below the 660 reflection gap (rather than along the slab dip), the responsible
velocity/density structure could have undergone retrograde motion during its descend into the
lower mantle. These observations collectively defines the large difference between Kuril and
Honshu slabs in terms of maximum vertical extension.
4.6. Other HRZs and potential inferences
Two additional anomalous reflectivity structures from the SS migration images could have
significant implications for the mantle structure, dynamics and/or mineralogy if confirmed.
First, we identify one (Japan subduction zone) or multiple (Kuril region) mid-MTZ HRZ(s) with
reflection amplitudes of 5-9% within the MTZ (see Fig. 3B and Fig. 4). With an exception of
one instance east from the slab (see Fig. 4, Profile A), these HRZs are consistently detected
within the slab contours suggested by Obayashi et al. (2006). Reflective structures near 520-km
depth have been documented nearly 3 decades ago in the Pacific northwest from travel time
observations (Fukao et al., 1977). It was later proposed to be a mild global seismic discontinuity
based on pioneering studies of SS precursors (Shearer, 1990, 1991). Bock (1994) explained this
reflector as a potential data processing artifact due to strong low-frequency side-lobes of S410S
and S660S phases, though more recent results based on reflected and converted body waves
(Gossler and Kind, 1996; Shearer, 1996; Flanagan and Shearer, 1998; Gu et al. 1998, Chevrot et
al., 1999; Deuss and Woodhouse, 2001; Gu et al., 2003; Lawrence and Shearer, 2006, Deuss,
2009) have favored an explanation that involves regionally variable, highly undulating reflective
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structure(s) in the MTZ. In terms of mineral physics, this interface has been attributed to
wadsleyite to ringwoodite (Helffrich, 2000, Bina, 2003) and/or garnet to Ca-perovskite (Ita and
Stixrude, 1992) phase transitions. In cold mantle regions such as subduction zones, these
transformations likely occur at different MTZ depths (Saikia et al., 2008) and produce multiple
reflectors (Deuss and Woodhouse, 2001; Deuss, 2009). This may be the case for the observed
HRZs within the Kuril slab. Alternatively, delayed meta-stable olivine phase transition (Sung
and Burns, 1976; Iidaka and Suetsugu, 1992; Jiang et al., 2008, Bina and Kawakatsu; 2010) and
the presence of water within slabs are also viable source of enhanced reflections in active plate
convergence zones. Ultimately, an accurate interpretation of the anomalous HRZs within the
MTZ is predicated upon a greater consensus on the mantle condition surrounding slabs, for
example, the water content. In view of the apparent north-to-south difference between Japan (a
single 520 reflector) and Kuril (multiple reflectors) subduction zones, a combination of these
mechanisms may be needed to properly explain our observations in the Pacific Northwest.
Lastly, a narrow MTZ and a series of strong HRZs east of the Benioff-zone (see Figs. 4) both
suggest low MTZ temperatures. This interpretation is supported by findings in recent studies of
ScS reverberations (Revenaugh and Sipkin, 1994; Bagley et al. 2009), seismic tomography
(Obayashi et al., 2006; Huang and Zhao, 2006; Zhao and Ohtani, 2009), and electrical
conductivity (Ichiki et al., 2006). Furthermore, the strong reflection from these structures (8-
12% of SS) may not be sufficiently explained by a thermal origin alone. Compositional
variations associated with a hot mantle plume, which was once active during the past 130 Ma,
could provide the additional source material necessary to accomodate some of the strong
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reflections detected in the depth range of 250- 700 km (see Figs. 3 and 4) (Obayashi et al. 2006;
Honda et al., 2007; Bagley et al. 2009; Li and van der Hilst, 2010).
Conclusions
The dynamic processes beneath northwestern Pacific are only a microcosm of those beneath
many subduction systems globally. For this reason, inferences based on our high-resolution
reflectivity images could be potentially applicable to other regions with similar tectonic settings.
Based on the spatial correlation between reflectivity and seismic velocity, we conclude that the
origins of the majority of highly reflective zones are thermal, instead of compositional, in nature.
The combined reflectivity and velocity information enables us to detect and interpret the
geometry and strengths of major mantle heterogeneities in the approximate depth range of 300-
1000 km.
shows clear signs of bending within the MTZ, but the center of the stagnant section of the slab
appears to be deformed or folded, as suggested by an average or shallow 660. The depths of the
two MTZ bounding olivine phase boundaries are negatively correlated if slab dip is considered.
We also identify strong seismic reflector(s) within the slab body within the MTZ through out the
The Honshu slab does not appear to extend below the transition zone. negative overall
correlation between the depths of the two major olivine phase boundaries. However, localized
topography on the 660 within the presumed stagnant part of the slab suggests significant vertical
deformation near the base of the upper mantle. A single reflector is identified at the depth range
of 500-540 km, which could be associated with by changes in T . which causes strong negative
correlations of the olivine phase boundary but and Kuril slabs In particular, our analysis
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demonstrated that ‘gaps’ in the reflection amplitudes of the 410 and 660 are potentially
interconnected with anomalous lower mantle reflectors. Major mass/heat fluxes, large
topography on the base of upper mantle, and lower-mantle thermal/composition variations would
be expected at these locations. Intermittent reflections within the MTZ offer additional
information on the geometries and dynamics of stagnant slabs. In other words, a self-consistent
model of mantle processes beneath subduction zones is tenable from the presence, strengths, and
depths of mantle reflectors and their spatial correlations with seismic velocities.
From a technical standpoint, the results presented in this study provide a glimpse of the future for
regional-scale analysis based on intermediate-period SS precursors. Increasingly diverse
applications in recent years (e.g., Schmerr and Garnero, 2006, 2007; Houser et al., 2008;
Lawrence and Shearer, 2008; Gu et al., 2009; Rychert and Shearer, 2009; Heit et al., 2010; Cao
et al., 2010; Houser and Williams, 2010) have underlined the remarkable resolving power of this
data set, one that was traditionally tapped as a ‘low resolution’ constraint on mantle structure.
This trend will likely continue in the foreseeable future, especially in view of the growing
number of global seismic networks and applications of array methods.
Acknowledgement
We sincerely thank Suzan van der Lee for her constructive scientific input to this study. We are
grateful to Peter Shearer for his patience and professionalism in handling this manuscript. This
study also benefited from the helpful comments and suggestions from Nicholas Schmerr and an
anonymous reviewer, as well as from the technical assistance from the IRIS Data Management
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Center. This project is jointly funded by CFI, Alberta Innovates, Alberta Geological Survey,
National Science and Engineering Council (NSERC), and the University of Alberta.
References
Ai, Y., Zheng, T., 2003. The upper mantle discontinuity structure beneath eastern China.
Geophysical Research Letters 30, doi:10.1029/2003GL017678.
Akaogi, M., A. Tanaka, and, E. Ito, Garnet-ilmenite-perovskite transitions in the systemMg4Si4O12-Mg3Al2Si3O12 at high pressures and high temperatures: phase equilibria, calorimetryand implications for the mantle structure, Phys. Earth Planet. Inter., 132, 303-324, 2002.
Akaogi, M., Takayama, H. Kojitani, H. Kawaji, H., Atake, T., 2007. Low-temperature heat
capacities, entropies and enthalpies of Mg2SiO4 polymorphs, and and post-spinel
phase relations at high pressure. Physics and Chemistry of Minerals 34, 169-183.
Bagly, B., 1982. Geometry of subducted plates and island arcs viewed as a buckling problem. Geology 10, 629-632.
Bagley, B., Courtier, A. M., Revenaugh, J., 2009. Melting in the deep upper mantle oceanward of
the Honshu slab. Physics of the Earth and Planetary Interiors 175, 137–144.
Bercovici, D., and S. Karato (2003), Whole mantle convection and thetransition‐zone water filter, Nature, 425, 39–44.
Bassin C., Laske, G., Masters, G., 2000. The current limits of resolution for surface wave
tomography in North America. EOS Trans AGU 81, F897.
H.M. Benz, J.E. Vidale, Sharpness of upper-mantle discontinuitiesdetermined from high-frequency reflections, Nature 365(1993) 147–150.
1
2
3
4
5
6
7
89
1011
12
13
14
15
161718
19
20
212223
24
25
26
27282930
31
Bijwaard, H.,Spakman, W., Engdahl, E. R., 1998. Closing the gap between regional and global
travel time tomography. Journal of Geophysical Research 103, 30055–30078.
Billen, M. I., 2008. Modeling the dynamics of subducting slabs. Annual Review of Earth and
Planetary Sciences 36, 325-356.
Billen, M. I., Hirth, G., 2007. Rheologic controls on slab dynamics. Geophysics, Geochemistry,
Geosystems 8, Q08012, doi:10.1029/2007GC001597.
Bina, C.R., Stein, S., Marton, F.C., Van Ark, E.M., 2001. Implications of slab mineralogyfor subduction dynamics. Phys. Earth Planet. Inter. 127, 51–66.
Bina, C. R., 2003. Seismological Constraints upon Mantle Composition, (in) Carson R. (Ed),
Treatise on Geochemistry vol 2., Elsevier Science Publishing, Oxford, pp. 39-59.
Bina, C. R., Kawakatsu, H., 2010. Buoyancy, bending and seismic visibility in deep slab
stagnation. Physics of the Earth and Planetary Interiors, 183, 330-340.
Bird, P., 2003. An updated digital model of plate boundaries. Geochemistry, Geophysics,
Geosystems 4, 1027.
Cao, Q., Van der Hilst, R.D., De Hoop, M.V., Shim, S.-H., 2011. Seismic imaging of transition
zone discontinuities suggests hot mantle west of Hawaii. Science 332, 1068-1071,
doi:10.1126/science.1202731.
Chaljub, E., Tarantola, A., 1997. Sensitivity of SS precursors to topography on the upper-mantle 660-km discontinuity. Geophysical Research Letters 24, 2613-2616.
Chang, S-J., van der Lee, S., Flanagan, M. P., Bedle, H., Marone, F., Matzel, E. M., Pasyanos, M.
E., Rodgers, A. J., Romanowicz, B., Schmid, C., 2010. Joint inversion for 3-dimensional S-
velocity mantle structure along Tethyan margin. Journal of Geophysical Research 115, B08309.
Chen, P., Bina, C.R., Okal, E.A., 2004. A global survey of stress orientations in subductingslabs as revealed by intermediate-depth earthquakes. Geophys. J. Int. 159,
1
2
3
4
5
6
7
8
91011
12
13
14
15
16
17
18
19
20
21
22
23
24
25
262728
29
30
31
323334
721–733.
Chen, L., and Y. Ai (2009), Discontinuity structure of the mantle transition zone
beneath the North China Craton from receiver function migration, J. Geophys.
Res., 114, B06307, doi:10.1029/2008JB006221.
Christensen, U., 1995. Effects of phase transitions on mantle convection. Annu. Rev. Earth
Planet. Sci. 23, 65–87.
Christensen UR. 1996. The influence of trench migration on slab penetration into the lower mantle. EarthPlanet. Sci. Lett. 140:27–39
Christensen, U., 2010. Geodynamic models of deep subduction. Phys. Earth Planet.Inter. 127, 25–34.
Collier, J., Helffrich, G., Wood, B., 2001. Seismic discontinuities in subduction zones.Phys. Earth Planet. Int. 127, 39–49.
Courtier, A. M., Revenaugh, J., 2006. A water-rich transition zone beneath the eastern United
States and Gulf of Mexico from multiple ScS reverberations. Earth’s Deep Water Cycle. Ed. S.D.
Jacobsen and S. van der Lee. Geophysical Monograph (AGU), 181–193.
Courtier, A. M., Revenaugh, J., 2008. Slabs and shear wave reflectors in the mid-mantle. Journal
of Geophysical Research 113, B08312.
DeMets, C., Gordon, R.G., Argus, D.F., Stein, S., 1990. Current plate motions. GeophysicalJournal International 101, 425–478.
Deuss, A. and Woodhouse, J. H., 2002. A systematic search for mantle discontinuities using SS-
precursors. Geophysical Research Letters, 29, 1249-1252.
1
2
3
4
5
6
7
8
91011
121314
15
161718
19
20
21
22
23
24
25
26
27
282930
31
32
33
34
Deuss, A., 2007. Seismic observations of transition zone discontinuities beneath
hotspot locations, Special Papers of the Geological Society of America, 430
(Plates, plumes and Planetary Processes edited by Foulger, G. R. and D. M.
Jurdy), 121-136.
Deuss, A., 2009. Global observations of mantle discontinuities using SS and PP precursors.
Surveys in Geophysics, 30 (4-5), 301-326.
Du, Z., Vinnik, L.P. & Foulger, G.R., 2006. Evidence from P-to-S mantleconverted waves for a flat ‘660-km’ discontinuity beneath Iceland, Earthplanet. Sci. Lett., 241, 271–280.
Dziewonski, A. M., Anderson, D. L., 1981. Preliminary reference earth model, Physics of the
Earth and Planetary Interiors 25, 297-356.
Dziewonski, A. and Woodhouse, J. (1983). Studies of the seismic source using normal-mode
theory. In Kanamori, H. and Boschi, E., editors, Earthquakes: observation, theory, and
interpretation: notes from the International School of Physics ``Enrico Fermi'' (1982: Varenna,
Italy), pages 45-137. North-Holland Publ. Co., Amsterdam.
Efron, B., 1977. Bootstrap method, another look at the Jackknife. The Annals of Statistics 7, 1-
26.
Fee, D. & Dueker, K., 2004. Mantle transition zone topography and structurebeneath the Yellowstone hotspot, Geophys. Res. Lett., 31, L18603,doi:10.1029/2004GL020636.
Flanagan, M.P., Shearer, P. M., 1998. Global mapping of topography on transition zone velocity
discontinuities by stacking SS precursors. Journal of Geophysical Research 103, 2673–2692.
Frost, D. 2008. The upper mantle and transition zone. Elements 4, 171-176.
Frost, D., Dolejs, D., 2007. Experimental determination of the effect of H2O on the 410-km
seismic discontinuity. Earth and Planetary Science Letters, 256, 182-195.
1
2
3
4
5
6
7
89
1011
12
13
14
15
16
17
18
19
20
21
22
23242526
27
28
29
30
31
32
33
34
Fuchs, K., and G. Muller, Computation of synthetic seismograms with the reflectivity method and comparison with observations, Geophys. J. R. Astr. Soc., 23, 417-433, 1971.
Fukao, Y., 1977. Upper mantle P structure on the ocean side of the Japan-Kuril Arc. Geophysical
Journal of Royal Astronomical Society 50, 621-642.
Fukao, Y., Obayashi, M., 2010. Transition from slab stagnation to penetration beneath the
northwestern Pacific and South America. American Geophysical Union, Fall Meeting, abstract
#DI23C-01.
Fukao, Y., Obayashi, M., Inoue, H., Nenbai, M., 1992. Subducting slabs stagnant in the Mantle
Transition Zone. Journal of Geophysical Research 97, 4809–4822.
Fukao, Y.,, Obayashi, M., Nakakuki, T., Deep Slab Project Group, 2009. Stagnant slab: A review.
Annual Review of Earth and Planetary Sciences 37, 19-46.
Fukao, Y., Widiyantoro, S., Obayashi, M., 2001. Stagnant slabs in the upper and lower mantle
transition region. Reviews of Geophysics 39, 291–323.
Gilbert, J.H., Sheehan, A.F., Dueker, K.G. & Molnar, P., 2003. Receiverfunctions in the western United States, with implications for uppermantle structure and dynamics, J. geophys. Res., 108(B5), 2229,doi:10.1029/2001JB001194.Gossler, J. & Kind, R., 1996.
Gasparik, T., 1996. Melting experiments on the enstatite-diopsidejoin at 70–224 kbar, including the melting of diopside. Contrib.Mineral. Petrol. 124, 139–153.
Gorbatov A, Kennett BLN. 2002. Joint bulk-sound and shear tomography for western Pacific subductionzones. Earth Planet. Sci. Lett. 210:527–43
Gorbatov A,Widiyantoro S, Fukao Y, Gordeev E. 2000. Signature of remnant slabs in the North Pacific fromP-wave tomography. Geophys. J. Int. 142:27–36
Gossler J., Kind, R., 1996. Seismic evidence for very deep roots of continents. Earth and
Planetary Science Letters 138, 1-13.
1
234
5
6
7
8
9
10
11
12
13
14
15
16
17
18
192021222324
25262728
293031
323334
35
36
37
Gu Y. J., Dziewonski, A. M., Agee, C. B., 1998. Global de-correlation of the topography of
transition zone discontinuities. Earth and Planetary Science Letters 157, 57-67.
Gu Y. J., Dziewonski, A. M., 2002. Global variability of transition zone thickness. Journal of
Geophysical Research 107, 2135.
Gu, Y. J., Dziewonski, A. M., Ekström, G., 2003. Simultaneous inversion for mantle shear
velocity and topography of transition zone discontinuities. Geophysical Journal International
154, 559–583.
Gu, Y. J., Schultz, R., Okeler, A., 2008. Migration and radon imaging of the western Pacific
subduction zones using SdS waves. EOS Trans AGU. 89, DI12A-08.
Gudmundsson O., Sambridge, M., 1998. A regionalized upper mantle (RUM) seismic model.
Journal of Geophysical Research 103, 7121-7136.
Heit, B., Yuan, X., Bianchi, M., Kind, R. And Gossler, J., 2010. Study of the lithospheric and
upper-mantle discontinuities beneath eastern Asia by SS precursors. Geophysical Journal
International 183, 252-266.
Helffrich, G., 2000. Topography of the transition zone seismic discontinuities. Reviews of
Geophysics 38, 141–158.
Herrmann, R. B., and C. Y. Wang, A comparison of synthetic seismograms, Bull. Seism. Soc. Am., 75, 41-
56, 1985.
Hirose K, Fei Y, Ma Y, Mao HK. 1999. The fate of subducted basaltic crust in the Earth’s lower mantle.Nature 397:53–56
Hirose K, Takafuji N, Sata N, Ohishi Y. 2005. Phase transition and density of subducted MORB crust in thelower mantle. Earth Planet. Sci. Lett. 237:239–51
1
2
3
4
5
6
7
8
9
10
11
12
13
14
15
16
17
18
19
20
21
22
23
24
25
26
272829
303132
33
Honda, S., Morishige, M., Orihashi, Y., 2007. Sinking hot anomaly trapped at the 410 km
discontinuity near the Honshu subduction zone, Japan. Earth and Planetary Science Letters 261,
565–577.
Houser, C., Williams, Q., 2010. Reconcilling Pacific 410 and 660 km discontinuity topography,
transition zone shear velocity patterns, and mantle phase transitions. Earth and Planetary Science
Letters 296, 255-266.
Houser, C., Masters, G., Flanagan, M., Shearer, P., 2008. Determination and analysis of long-
wavelength transition zone structure using SS precursors, Geophysical Journal International 174,
178-194.
Huang, J., Zhao, D., 2006. High-resolution mantle tomography of China and surrounding
regions. Journal of Geophysical Research 111, B09305.
Ichiki, M., Baba, K., Obayashi, M., Utada, H., 2006. Water content and geotherm in the upper
mantle above the stagnant slab: Interpretation of electrical conductivity and seismic P-wave
velocity models. Physics of The Earth and Planetary Interiors 155, 1 – 15.
Iidaka, T., Suetsugu, D., 1992. Seismological evidence for metastable olivine inside a subducting
slab. Nature 356, 593–595.
T. Inoue, H. Yurimoto, Y. Kudoh, Hydrous modified spinel,Mg1.75SiH0.5O4: a new water reservoir in the mantle transitionregion, Geophys. Res. Lett. 22 (1995) 117–120.
Inoue, T., Ueda, T., Tanimoto, Y., Yamada, A., Irifune, T., 2010. The effect of water on the high-
pressure phase boundaries in the system Mg2SiO4-Fe2SiO4. Journal of Physics: Conference Series
215, doi:10.1088/1742-6596/215/1/012101.
Irifune T, Ringwood AE. 1993. Phase transformations in subducted oceanic crust and buoyancy relationshipsat depths of 600–800 km in the mantle. Earth Planet. Sci. Lett. 117:101–10
1
2
3
4
5
6
7
8
9
10
11
12
13
14
15
16
17
18
19
20
21
22
23242526
27
28
29
30
313233
Ita, J., Stixrude, S., 1992. Petrology, elasticity, and composition of the mantle transition zone.
Journal of Geophysical Research 97, 6849–6866.
Irifune, T., Nishiyama, N., Kuroda, K., Inoue, T., Isshiki, M., Utsumi, W., Funakoshi, K.,Urakawa, S., Uchica, T., Katsura, T., Ohtaka, O., 1998. The postspinel phase boundary inMg2SiO4 determined by in situ X-ray diffraction. Science 279, 1698–1700.
Ji, Y., Nataf, H. C. 1998. Detection of mantle plumes in the lower mantle by diffraction tomography: Hawaii. Earth and Planetary Science Letters 159, 99-115.
Jasbinsek, J. J., Dueker, K. G., Hansen, S. M., 2010. Characterizing the 410 km discontinuity
low-velocity layer beneath the LA RISTRA array in the North American Southwest.
Geochemistry, Geophysics, Geosystems 11, Q03008, doi:10.1029/2009GC002836.
Jiang, G., Zhao, D., Zhang, G., 2008. Seismic evidence for a metastable olivine wedge in the
subducting Pacific slab under Japan Sea. Earth and Planetary Science Letters 270, 300–307.
Jordan, T. H. (1977). Lithospheric slab penetration into the lower mantle beneath the Sea of Okhotsk. Journal of Geophysics - Zeitschrift Fur Geophysik. Vol. 43 (1-2), pp. 473-496.
Kaneshima S, Helffrich G. 1999. Dipping low-velocity layer in the midlower mantle: evidence for geochemicalheterogeneity. Science 283:1888–91.
S. Karato, Remote sensing of hydrogen in the Earth's mantle,Rev. Mineral. Geochem. 62 (2006) 343–375.
Karato S,Wu P. 1993. Rheology of the upper mantle: a synthesis. Science 260:771–78
Katsura, T., Ito, E., 1989. The system Mg2SiO4-Fe2SiO4 at high pressures and temperatures; precise determination of stabilities of olivine, modified spinel, and spinel. Journal of Geophysical Research 94, 15663-15670.
Kawakatsu, H., Niu, F., 1994. Seismic evidence for a 920-km discontinuity in the mantle. Nature
371, 301–305.
Kellogg, L. H., Hager, B. H., van der Hilst, R. D., 1999. Compositional stratification in the deep
mantle. Science 283, 1881-1884.
1
2
3
4567
8
9101112
13
14
15
16
17
18
19
202122
232425
262728
2930
31323334
35
36
37
38
Kind, R., The reflectivity method for a buried source, J. Geophys., 44, 603-612, 1978.
D.L. Kohlstedt, H. Keppler, D.C. Rubie, Solubility of water inthe α, β and γ phases of (Mg,Fe)2SiO4, Contrib. Mineral. Petrol.123 (1996) 345–357.
Kubo, T., Ohtani, E., Kondo, T., Kato, T., Toma, M., Hosoya, T., Sano, A., Kikegawa, T., Nagase,
T., 2002. Metastable garnet in oceanic crust at the top of the lower mantle. Nature 420, 803–806.
Lawrence, J.F., and M. E. Wysession, Seismic evidence for subduction-transported water in the
lower mantle, in Earth's Deep-Water Cycle. AGU Monograph, 251-261, 2006.
Lawrence, J.F., and P.M. Shearer, A global study of transition zone Thickness using receiver
functions, J. Geophys. Ress., 111, dio:10.1029/2005JB003973, 2006.
Lawrence, J.F., and P.M. Shearer, Imaging mantle transition zone thickness with SdS-SS finite-
frequency sensitivity kernels, Geophys. J. Int., 174, 143-158, doi: 10,1111/j.1365-
246X.2007.03673.x, 2008.
Leahy, G. M., and D. Bercovici (2007), On the dynamics of a hydrous meltlayer above the transition zone, J. Geophys. Res., 112, B07401,doi:10.1029/2006JB004631.
Leahy, G. M., and D. Bercovici (2010), Reactive infiltration of hydrous melt above the mantle transition zone, J. Geophys. Res., 115, B08406,doi:10.1029/2009JB006757.
Lebedev, S., Chevrot, S., van der Hilst, R. D., 2006. The 660-km discontinuity with the
subducting NW-Pacific lithospheric slab. Earth and Planetary Science Letters 205, 25-35.
Lei, J., Zhao, D., 2005. P-wave tomography and origin of the Changbai intraplatevolcano in Northeast Asia. Tectonophysics 397, 281–295.
Li, C., van der Hilst, R. D., 2010. Structure of the upper mantle and transition zone beneath
Southeast Asia from traveltime tomography. Journal of Geophysical Research 115, B07308.
1
23456
7
8
9
10
11
12
13
14
15
16
17
18
19
20212223
24252627
28
29
30
3132
33
34
35
36
Li, J., Chen, Q.-F., Vanacore, E, Niu, F., 2008. Topography of the 660-km discontinuity and
impliations for a retrograde motion. Geophysical Research Letters 35, L01302, doi:10.1029/
2007GL031658.
Li, X., Sobolev, S. V., Kind, R., Yuan, X., Estabrook, C. H., 2000. A detailed receiver function
image of the upper mantle discontinuities in the Japan subduction zone. Earth and Planetary
Science Letters 183, 527–541.
Litasov, K., and E. Ohtani (2002), Phase relations and melt compositions inCMAS‐pyrolite‐H2O system up to 25 GPa, Phys. Earth Planet. Inter.,134, 105–127.
Litasov KD, Ohtani E, Sano A. 2006. Influence of water on major phase transitions in the Earth’s mantle. SeeJacobsen & van der Lee 2006, pp. 95–111
Mégnin, C., Romanowicz, B., 2000. The 3D shear velocity structure of the mantle from the
inversion of body, surface and higher mode waveforms. Geophysical Journal International 143,
709-728.
Melbourne, T., Helmberger, D., 1998. Fine structure of the 410-km discontinuity. Journal of
Geophysical Research 103, 10091-10102.
F. Neele, Sharp 400-km discontinuity from short-period Preflections, Geophys. Res. Lett. 23 (1996) 419–422.
Neele, F., de Regt, H.,Van Decar, J., 1997. Gross errors in upper-mantle discontinuity topography from underside reflection data, Geophysical Journal International 129, 194-204.
Niu, F., Levander, A., Ham, S., Obayashi, M., 2005. Mapping the subducting Pacific slabbeneath southwest Japan with Hi-net receiver functions. Earth and Planetary Science Letters 239, 9–17.
Northrup, C. J., Royden, L. H., Burchfiel, B. C., 1995. Motion of the Pacific plate relative to
Eurasia and its potential relation to Cenozoic extension along the eastern margin of Eurasia.
Geology 23, 719–722.
1
2
3
4
5
6
7
89
1011
121314
15
16
17
18
19
20
21
222324
25
26272829303132
33
34
35
36
Obayashi, M., Sugioka, S., Yoshimitsu, J., Fukao, Y., 2006. High temperature anomalies
oceanward of subducting slabs at the 410-km discontinuity. Earth and Planetary Science Letters
243, 149–158.
Ohtani, E., 2005. Water in the Mantle. Elements 1, 25–30.
Okino K, Ando M, Kaneshima S, Hirahara K. 1989. The horizontally lying slab. Geophys. Res. Lett. 16:1059–2062
Press, W. H., Flannery, B. P., Teukolsky, S. A., Vetterling, W. T.. 1992. Numerical recipes in C:
The Art of scientific computing. 2nd Edition, Cambridge University Press, Cambridge, United
Kingdom.
Ramesh, D., Kawakatsu, H., Watada, S., Yuan, X., 2005. Receiver function images of thecentral Chugoku region in the Japanese islands using Hi-net data. Earth PlanetsSpace 57, 271–280.
Revenaugh, J., Jordan, T. H., 1991. Mantle layering from ScS reverberations: 2. the transition
zone. Journal of Geophysical Research 96, 19763-19810.
Revenaugh, J., Sipkin, S. A., 1994. Seismic evidence for silicate melt atop the 410-km mantle
discontinuity. Nature 369, 474–476.
Richard, G., Bercovici, D., Karato, S -I., 2006. Slab dehydration in the Earth’s mantle transition
zone. Earth and Planetary Science Letters 251, 156–167.
Rondenay, S., 2009. Upper mantle imaging with array recordings of converted and scattered
teleseismic waves. Survey of Geophysics 30, 377.
Rychert, C. A., Shearer, P. M., 2009. A global view of the lithosphere-asthenosphere boundary.
Science 332, 495-498, doi:10.1126/science.1169754.
Saita, T., Suetsugu, D., Ohtaki, T., Takenaka, H., Kanjo, K., Purwana, I., 2002. Transitionzone thickness beneath Indonesia as inferred using the receiver function methodfor data from JISNET regional broadband seismic network. Geophys. Res. Lett. 29.doi:10.1029/2001GL013629.
1
2
3
4
5
678
9
10
11
12
13141516
17
18
19
20
21
22
23
24
25
26
27
28
29
30
31
3233343536
Schaeffer, A. J., Bostock, M. G., 2010. A low velocity zone atop the transition zone in
northwestern Canada. Journal of Geophysical Research 15, B06302,
doi:10.1029/2009JB006856.
Schmerr, N, Garnero, E. J., 2006. Investigation of upper mantle discontinuity structure beneath the central Pacific using SS precursors. Journal of Geophysical Research 111, B08305, doi:10.1029/2005JB004197.
Schmerr, N., Garnero, E. J., 2007. Upper mantle discontinuity topography from thermal and
chemical heterogeneity. Science 318, 623-626.
Schmid C, Goes S, van der Lee S, Giardini D. 2002. Fate of the Cenozoic Farallon slab from a comparison ofkinematic thermal modeling with tomographic images. Earth Planet. Sci. Lett. 204:17–32
Seno, T., Sakurai, T., Stein, S., 1996. Can the Okhotsk plate be discriminated from the NorthAmerican plate? Journal of Geophysical Research 101, 11305–11315.
Shearer, P. M., 1993. Global mapping of upper mantle reflectors from long-period SS precursors.
Geophysical Journal International 115, 878-904.
Shearer, P. M., 1996. Transition zone velocity gradients and the 520-km discontinuity. Journal of
Geophysical Research 101, 3053-3066.
Shearer, P. M., 2000. Upper mantle seismic discontinuities. Earth’s Deep Interior: Mineral
Physics and Tomography from the Atomic to the Global Scale, Geophysical Monograph 117,
115-131.
Shearer, P. M., Masters, T. G., 1992. Global mapping of topography on the 660 km discontinuity.
Nature 355, 791–796.
1
2
34
5
6789
10
11
121314
15
16171819
20
21
22
23
24
25
26
27
28
29
30
31
32
33
Shearer, P. M., Flanagan, M. F., Hedlin, A. H., 1999. Experiments in migration processing of SS precursor data to image upper mantle discontinuity structure. Journal of Geophysical Research, 104, 7229-7242.
Shen, X., Zhou, H., Kawakatsu, H., 2008. Mapping the upper mantle discontinuities beneath
China with teleseismic receiver functions. Earth Planets Space 60, 713–719.
Shieh, S. R., Mao, H. -K, Hemley, R. J., Chung Ming, L., 1998. Decomposition of phase D in
the lower mantle and the fate of dense hydrous silicates in subducting slabs. Earth and Planetary
Science Letters 159, 13–23.
Smyth, J. R., Frost, D. J., 2001. The effect of water on the 410-km discontinuity: An
experimental study. Geophysical Research Letters 29, doi:10.1029/2001GL014418.
R. Smyth, S. D. Jacobsen, in Earth's Deep Water Cycle, S. D. Jacobsen, S. Van der Lee, Eds. (American Geophysical Union, Washington, DC, 2006), 168, 1–11. 32.
Stixrude, L., Lithgow-Bertelloni, C., Kiefer, B., Fumagalli, P., 2007. Phase stability and shear
softening in CaSiO3 perovskite at high pressure. Phys. Rev. B. 75, 024108.
Suetsugu, D., Inoue, T., Yamada, A., Zhao, D., Obayashi, M., 2006. Towards mappingthe three-dimensional distribution of water in the transition zone fromP-velocity tomography and 660-km discontinuity depths, geophysical monographseries 167, ‘Earth’s deep water cycle’. AGU, 237–249.
Suetsugu, D., Inoue, T., Obayashi, M., Yamada, A., Shiobara, H., Sugioka, H., Ito, A., Kanazawa,
T., Kawakatsu, H., Shito, A., Fukao, Y., 2010. Depths of the 410-km and 660-km discontinuiies
in and around the stagnant slab beneath the Philippine Sea: Is water stored in the stagnant slab?
Physics of the Earth and Planetary Interiors, 183, 270-279.Sugioka, H., Suetsugu, D., Obayashi, M., Fukao, Y., Gao, Y., 2010. Fast P and S wavevelocities associated with the “cold” stagnant slab beneath the northern PhilippineSea. Phys. Earth Planet. Inter. 179, 1–6, doi:10.1016/j.pepi.2010.01.006.
Sung, C. -M., Burns, R. G., 1976. Kinetics of high-pressure phase transformations: Implications
to the evolution of the olivine-spinel transition in the downgoing lithosphere and its
consequences on the dynamics of the mantle. Tectonophysics 31, 1–32.
Tagawa M, Nakakuki T, Tajima F. 2007. Dynamical modeling of trench retreat driven by the slab interaction
1234
5
6
7
8
9
10
11
12
13
141516
171819
20
2122232425
26
27
28
29
30313233
34
35
36
37
3839
with the mantle transition zone. Earth Planets Space 59:65–74
B. Tauzin, E. Debayle, G. Wittlinger (2008), The mantle transition zone as seen by global Pds
phases : no clear evidence for a thin transition zone beneath hotspots, J. Geophys. Res., 113,
B08309, doi:10.1029/2007JB005364.
B. Tauzin, E. Debayle, G. Wittlinger (2010), Seismic evidence for a global low velocity layer in the Earth's upper mantle, Nature Geoscience, 3, 718-721, doi:10.1038/NGEO969.
Tetzlaff, M., Schmeling, H., 2000. The influence of olivine metastability ondeep subduction of oceanic lithosphere. Phys. Earth Planet. Inter. 120,29–38.
Tibi, R., Wiens, D. A., Shiobara, H., Sugioka, H., Yuan, X., 2007. Geophysical Research Letters
34 ,L16316, doi:10.1029/2007GL030527.
Tonegawa, T., Hirahara, K., Shibutani, T., Fujii, N., 2006. Lower slab boundary in the Japan
subduction zone. Earth and Planetary Science Letters 247, 101 – 107.
Torii, Y., Yoshioka, S., 2007. Physical conditions producing slab stagnation: constraintsof the Clapeyron slope, mantle viscosity, trench retreat, and dip angles.Tectonophysics 445, 200–209.
Vacher, P., A. Mocquet, and C. Sotin, Computation of seismic profiles from mineral physics: theimportance of the non-olivine components for explaining the 660 km depth discontinuity, Phys.Earth Planet. Inter., 106, 275-298, 1998.
van der Meijde, M., Marone, F., Giardini, D., van der Lee, S., 2003. Seismic evidence for water
deep in Earth’s upper mantle. Science 300, 1556–1558.
van der Meijde, M., van der Lee, S., Giardini, D., 2005. Seismic discontinuities in the
Mediterranean mantle. Physics of the Earth and Planetary Interiors 148, 233–250.
van der Hilst, R., Engdahl, R., Spakman, W., Nolet, G., 1991. Tomographic imaging of
subducted lithosphere below northwest Pacific island arcs. Nature 353, 37–43.
1
2345
6
789
101112
13
14
15
16
17
18
19202122
232425262728
29
30
31
32
33
34
35
36
37
van der Hilst, R. D., Widiyantoro, S., Engdahl, E. R., 1997. Evidence for deep mantle circulation
from global tomography. Nature 386, 578–584.
van Keken PE, Karato S, Yuen DA. 1996. Rheological control of oceanic crust separation in the transitionzone. Geophys. Res. Lett. 23:1821–24.
Vidale, J. E., Ding, X.-Y., Grand, S. P., 1995. The 410-km-depth discontinuity: A sharp estimate
from near-critical reflections. Geophysical Research Letters 22, 2557-2560.
Wang, Y., Uchida, T., Zhang, J., Rivers, M. L., Sutton, S. R., 2004. Thermal equation of state of
akimotite MgSiO3 and effect of akimotite-garnet transformation on seismic structure near the
660 km discontinuity. Physics of the Earth and Planetary Interiors 143-144, 57-80.
Weidner, D., Wang, Y., 1998. Chemical- and Clapeyron-induced buoyancy at the 660 kmdiscontinuity. J. Geophys. Res. 103, 7431–7441.
Weidner, D., Wang, Y., 2000. Phase transformations; implications for mantle structure.In: Karato, S., Forte, A., Liebermann, R., Masters, G., Stixrude, L (Eds.), Earth's DeepInterior: Mineral Physics and Tomography from the Atomic to Global Scale.American Geophysical Union, pp. 215–235.
Wen, L., Anderson, D. L., 1997. Layered mantle convection: A model for geoid and topography.
Earth and Planetary Science Letters 146, 367 – 377.
Widiyantoro, S., Kennett, B.L.N., van der Hilst, R.D., 1999. Seismic tomography withP and S data reveals lateral variations in the rigidity of deep slabs. Earth Planet.Sci. Lett. 173, 91–100.
B.J. Wood, The effect of H2O on the 410-kilometer seismicdiscontinuity, Science 268 (1995) 74–76.
Yamada, A., Zhao, D., Inoue, T., Suetsugu, D., Obayashi, M., 2009. Seismological evidence for
compositional variations at the base of the mantle transition zone under Japan islands. Gondwana
Research, doi:10.1016/j.gr.2009.04.009
1
2
3
456
7
8
9
10
11
12
13
14
15161718192021222324
25
26
27
28293031
323334
35
36
37
38
Zhao, D., 2004. Global tomographic images of mantle plumes and subductingslabs: insight into deep Earth dynamics. Phys. Earth Planet. Inter. 146, 3–34.Zhao, D., Hasegawa, A., Horiuchi, S., 1992. Tomographic imaging of P and Swave velocity structure beneath northeastern Japan. J. Geophys. Res. 97,19,909–19,928.
Zhao, D., Ohtani, E., 2009. Deep slab subduction and dehydration and their geodynamic
consequences: Evidence from seismology and mineral physics. Gondwana Research 16, 401–
413.
Zhu, G., Shi, Y., Tackley, P., 2010. Subduction of the Western Pacific Plate underneath Northeast
China: Implications of numerical studies. Physics of the Earth and Planetary Interiors 178, 92-99.
Mantle dynamics of Western Pacific and East Asia: Insight fromseismic tomography and mineral physicsDapeng Zhao a, Shigenori Maruyama b, Soichi Omori, Gondwana Research 11 (2007) 120–131.
Zhu, G., Shi, Y., Tackley, P., 2010. Subduction of the Western Pacific Plate underneath Northeast
China: Implications of numerical studies. Physics of the Earth and Planetary Interiors 178, 92 –
99.
Figure Captions:
Fig. 1: (A) A schematic drawing of SS precursor reflection from a subducting oceanic lithosphere
at the base of upper mantle. These waves are sensitive to the depth and impedance contrast of a
mantle interface. (B) Ray theoretical surface reflection points of 6014 high-quality SS waves
used in this study. The main tectonic elements and plate boundaries (Bird, 2003) and slab
contours (Gudmundsson and Smabridge, 1998) are shown by thick and thin black lines,
respectively. The surface projections of five mantle transects (see main text) are labeled A-E,
extending from central Honshu Island (A) to central Kuril Arc (E).
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Fig. 2: Key steps in the time-to-depth migration of SS precursors. By placing the aligned SS
precursors at the surface (Middle), time samples of transversely polarized seismograms prior to
the arrival of SS (Left) can be effectively mapped to corresponding reflection depths (2nd to the
Right) along the predicted differential time curves based on PREM (Dziewonski and Anderson,
1981). The Right-most panel shows the isotropic shear velocities of PREM down to 1800-km
depth.
Fig. 3: Interpolated reflectivity maps of SS precursor amplitude variations at MTZ (A to C) and
(D) shallow lower mantle depths. An isosurface (threshold = 7.5%) is used to define HRZ in all
panels.. The anomalies marked with green dashed lines are discussed in the text. Slab depth
contours (Gudmundsson and Sambridge, 1998) are drawn by magenta lines at 50 km intervals
from the trench.
Fig. 4: Interpolated CMP gathers along profiles A to D (see Fig. 1A) superimposed on high-
resolution P-wave velocities (Obayashi et al., 2006). Also indicated are earthquakes within the
averaging window of each cross-section. Our interpretations (white lines) are combined with the
-0.5% velocity perturbation contours (red lines, Obayashi et al., 2006).
Fig. 5: (A) Mantle reflectivity structures along the northernmost profile E (see Fig. 1B), P-wave
speeds (Obayashi et al., 2006), and Wadati-Benioff zone seismicity (yellow circles). The thin
red-lines outlines -0.5% velocity perturbations (Obayashi et al., 2006). Our interpretations are
highlighted by the dashed white lines. (B) A schematic interpretation of the HRZs for the Japan-
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Kuril subduction system. The thick red line along the surface of the subducting slab indicates the
ongoing process of dehydration melting. Slab penetration is likely in regions where the 660
appears to be segmented.
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