Tracking Slabs Beneath Northeastern Pacific Subduction Zones

2 Tracking Slabs Beneath Northeastern Pacific Subduction Zones

4 Yu Jeffrey Gu

5 University of Alberta, Department of Physics, CEB 348-D, Edmonton, AB, Canada, T6G 2G7.

6 E-mail: [email protected]

7 Phone: 1 780 492 2292

8 Fax: 1 780 492 0714

10 Ahmet Okeler

11 University of Alberta, Department of Physics, CEB 456, Edmonton, AB, Canada, T6G 2G7.

13 Phone: 1 780 492 4125

14 Fax: 1 780 492 0714

16 Ryan Schultz

17 University of Alberta, Department of Physics, CEB 456, Edmonton, AB, Canada, T6G 2G7.

19 Phone: 1 780 492 4125

20 Fax: 1 780 492 0714

2 Abstract

4 Illuminating major thermal and/or compositional variations in Earth's mantle based on reflected

5 seismic waves is analogous to “motion tracking” in animation cinematography. Signals analyzed

6 by both approaches are sensitive to strong gradients in material properties and, with proper

7 treatments, can be used to decipher the shape or movements of the enclosed mass. In the same

8 spirit, this study utilizes the amplitudes of bottom-side reflected shear waves to provide first-

9 order constraints on the geometry and kinematics of subducted oceanic crust and lithosphere

10 beneath the northwestern Pacific subduction zones. The high-resolution, depth-migrated

11 reflection amplitudes shows large, ~1000 km wide depressions on the 660-km seismic

12 discontinuity, extending from the Japan sea to eastern China. The 410-km seismic discontinuity

13 is locally elevated by ~15 km on the oceanside of the Japan trench, where a sharp change of

14 transition zone thickness infers a mantle temperature increase over XX deg C. The 410-km

15 seismic discontinuity is locally elevated by ~15 km east of the Wadati-Benioff zone, within

16 which reflection amplitude drops off significantly. We further identify a strong reflector at ~530

17 km depth with a reflection amplitude exceeding 5% of SS amplitude. The strength of this

18 anomaly increases depressed with ‘avalanching’ the lower mantle west of the Hokkaido corner.

19 Strong correlations between the reflectivity structure and seismic velocity suggest: (1) high-

20 amplitude reflections generally occurs near the edges of major seismic anomalies due to strong

21 shear wave focusing effect, (2) ‘gaps’ in the reflection amplitudes of the 410- and 660-km

22 seismic discontinuities are associated with substantial topography and major mass/heat fluxes.,

2 1 and (3). The presence this reflectors residual plume(s) in this region. UNFINISHED, will work

2 on last.

4 1. Introduction

5 The convergent boundary between the Pacific, Amurian, and North American plates represents

6 one of the fastest destruction zones of old oceanic domains. The subduction process in this re-

7 gion initiated during the Cretaceous times (~65-140 Ma ago) (Northrup et al., 1995; Tonegawa et

8 al., 2006; Zhu et al., 2010) and continues to accommodate the differential motions between the

9 Pacific, Eurasia, and North American plates. The deposition of old oceanic lithosphere at the

10 present rate of 8-9.5 cm/yr (DeMets et al, 1990; Seno et al., 1996; Bird, 2003) not only directly

11 influences the surrounding mantle temperature and/or mineralogy.

13 The morphology and spatial extent of subducted oceanic lithosphere (for short, ‘slab’) beneath

14 the northwestern Pacific margin have long been investigated. Among the various data types and

15 approaches, seismic tomography of body waves has been the most effective in constraining

16 details of slab geometry and surrounding mantle conditions in this region (e.g., van der Hilst et

17 al., 1991, 1997; Fukao, 1992; Bijwaard et al., 1998; Fukao et al., 2001; Obayashi et al. 2006;

18 Huang and Zhao, 2006; Zhao and Ohtani, 2009; Li and van der Hilst, 2010). Well-defined zones

19 of above-average P and S wave speeds have been identified along the Wadati-Benioff zone and

20 within the upper mantle transition zone near Korea and eastern China (e.g., Jordan, 1977; van der

21 Hilst et al., 1997; Widiyantoro et al., 1997; Bijwaard et al., 1998; K´arason and van der Hilst,

22 2000; Fukao et al., 2001; Gorbatov et al., 2000; Gorbatov and Kennet, 2002; Lebedev and Nolet,

23 2003; Zhao, 2004; Obayashi et al., 2006; Huang and Zhao, 2006; Fukao et al., 2009; Zhao and 3 1 Ohtani, 2009; Li and van der Hilst, 2010). The non-geometrical shape of the high-velocity zones

2 have inspired discussions of slab deflection toward the horizontal, which is generally referred to

3 as ‘stagnation’ (Fukao et al., 1992; Fukao et al., 2001), and possible extension into the lower

4 mantle (see Fukao et al., 2001, 2009 for detailed reviews). The length of the flattened part of the

5 slab can be as large as 800-1000 km (Huang and Zhao, 2006; Obayashi et al., 2006; Fukao et al.,

6 2009), at least half of which can be reproduced numerically with proper treatments of trench

7 migration and rollback rates (). Low-velocity structures such as arc volcanism and/or

8 decompressional melting of stagnant slabs (Lebedev and Nolet, 2003; Zhao, 2004; Priestley et

9 al., 2006; Obayashi et al., 2006; Zhao and Ohtani, 2009; An et al., 2009; Wang et al., 2009; Duan

10 et al., 2009; Zhao et al., 2009; Li and van der Hilst, 2010; Feng and An, 2010), or hot thermal

11 plume(s) (Miyashiro, 1986; Ichiki et al., 2006; Zou et al., 2008; Zhao and Ohtani, 2009; Duan et

12 al., 2009), further underscores the wide range of dynamical processes beneath this region. These

13 low- and high-velocity heterogeneities can cause strong gradients in mantle temperature and/or

14 composition surrounding the convergent plate boundary zones

16 In comparison with seismic tomography, which is highly effective in resolving ‘smooth’

17 variations, the amplitudes and arrival-times of body waves reflected and converted at mantle

18 depths are more sensitive to sharp changes in rock elastic properties (Zheng et al., 2007).

19 Correlations between velocity and reflectivity (Shearer and Masters, 1992; Flanagan and Shearer,

20 1998; Li et al., 2000; Shen et al., 2008) offer greater constraints on slab geometry and dynamics

21 than either approach alone. For this reason, the temperature-dependent depressions on the 660-

22 km seismic discontinuity by 15-60 km (Shearer and Masters, 1992; Benz and Vidale, 1992; Bina

23 and Helfrich? Helfrich and Bina?? Li et al., 2000; Niu et al., 2005; Tonegawa et al., 2005; Shen

4 1 et al., 2008; Tauzin et al., 200??; Lawrence and Shearer, 2006; Houser et al., 2008) have been

2 widely cited as evidence of stagnating and ponding slab beneath the northwestern Pacific

3 collision zone. The resolutions of these seismic surveys are, however, hampered by the

4 restrictive source-receiver distributions of converted phases and the large averaging radii in

5 global analyses of secondary reflected waves also known as ‘SS precursors’. In particular, while

6 a pioneering study of the latter phase (Shearer and Masters, 1992) provided evidence of

7 stagnating slab beneath the northwestern Pacific region nearly 20 years ago, further usage of

8 these phases in constraining detailed slab geometry and kinematics was debated (Neele et al.,

9 1997; Shearer et al., 1999). Discussions of the correlations between mantle reflectivity inferred

10 from SS precursors and seismic velocities/mantle mineralogy near subduction zones mainly

11 focused on broad length scales and remained qualitative (e.g., Gu et al., 2003; Lawrence and

12 Shearer, 2006; Houser et al., 2008).

13 14 This study analyzes a large regional dataset of SS precursors using novel processing techniques

15 to improve the resolution on the seismic reflectivity structure beneath the northwestern Pacific

16 region (Fig. 1A). The dense regional data coverage enables pre-stack depth migration that

17 positions weak SS precursor amplitudes at the appropriate reflection depths and locations. By

18 correlating reflection amplitude variations with wave speeds, we aim to provide a self-consistent,

19 three-dimensional (3D) snapshot of mantle reflectivity structure and deformation near the

20 northwestern segment of the Pacific Ocean basin. For brevity we will hereon refer to the upper

21 mantle transition zone as MTZ and the 410-km, 520-km and 660-km discontinuities as the 410,

22 520 and 660, respectively.

23 5 1 2. Data and method

2 SS precursors are a proven means for determining the depths of mantle reflectors (e.g., Shearer

3 and Masters, 1992; Shearer, 1993; Gossler and Kind, 1996; Gu et al., 1998; Deuss and

4 Woodhouse, 2002; Flanagan and Shearer, 1998; Gu and Dziewonski, 2002; Schmerr and

5 Garnero, 2007; Lawrence and Shearer, 2007; Houser et al., 200XX; Rychert and Shearer, ??).

6 Their strong sensitivities to the reflection depth and interfacial impedance contrast beneath mid

7 points (see Fig. 1A), coupled with their strong sensitivity to structures away from the source and

8 station locations, are ideal for mapping mantle reflectivity at both global and regional scales.

10 We utilize all available broadband, high-gain recordings of earthquakes that took place prior to

11 2008. This data set is currently managed by the IRIS Data Management Center and highlights

12 significant efforts from GDSN, IRIS, GEOSCOPE and several other temporary deployments.

13 Only data from shallow events (<75 km) with magnitude (Mw) grater than 5.0 are selected for

14 this undertaking. The former criterion minimizes the effect of depth phase, and the subjective

15 magnitude cutoff ensure that source mechanism solutions are available from the Global Centroid

16 Moment Tensor (GCMT) project (Dziewonski and Woodhouse, 1983) for accurate computations

17 of PREM (Dziewonski and Anderson, 1981) synthetic seismograms. We further restrict the

18 epicenter distance range to 100°-160° to minimize known waveform interferences from topside

19 reflection sdsS and ScS precursors ScSdScS, where d denotes a discontinuity (Schmerr and

20 Garnero, 2007). After applying a Butterworth band-pass filter with corner periods at 12 s and 75

21 s to the selected data traces, we impose a signal-to-noise ratio (SNR) criterion as the ratio

22 between the maximum absolute amplitude of the SS and noise. The selected signal and noise

23 windows are (-20 sec, 60 sec) and (-170 sec, -80 sec), respectively, relative to the predicted

6 1 arrival time of SS based on PREM (Dziewonski and Anderson, 1981). All records with SNR

2 lower than 3.0 are automatically rejected.

4 The selected transverse-component seismograms are subsequently aligned on the first major

5 swing of SS phase with the aid of the corresponding synthetic seismograms. As the last step of

6 pre-processing, we apply time shifts by the theoretical SS and S520S times through PREM

7 (Dziewonski and Anderson, 1981) to account for crustal (Bassin et al., 2000) and topographical

8 (ETOPO5 data base) variations. Since our main focus is the upper mantle transition zone, the

9 approximation based on SS-S520S represents an effective compromise between the 410 and 660

10 and may introduce an error of 3-5 km for the depth estimation of reflectors hundreds of

11 kilometers away from the MTZ. Generally, these model assumptions have greater impacts on the

12 differential times, hence reflection depths, than on the amplitudes of SS precursors (e.g., Gu et

13 al., 2003).

15 A time-to-depth migration approach, which has been previously applied to P-to-S converted

16 waves (Rondenay, 2009 and references therein), is introduced to convert the precursory arrivals

17 of SS waves to the corresponding reflection depth and location (Gu et al., 2008; Heit et al.,

18 2010). The SS waveforms after the corrections for crust thickness and surface topography

19 correspond to equalized reflection at the Earth’s surface. Hence, each time sample preceding the

20 reference SS time can be mapped to a crustal/mantle depth according to the predicted travel-time

21 tables computed based on PREM (Dziewonski and Anderson, 1981) (Fig. 1B). The sampling

22 rate along the depth axis is 1 km.

23 7 1 Finally, to obtain a 3D reflectivity image we divide the study region into uniform, rectangular

2 Common Mid Point (CMP) gathers with horizontal and vertical step sizes of 2° and 8°,

3 respectively (IS THIS TRUE, AHMET?)_. Time-to-depth migration (Zheng et al., 2007) is

4 subsequently performed at each cell and the entire set of resulting migrated traces is interpolated

5 using a 3D, bi-linear interpolation method provided by MATLAB. Despite linear interpolation

6 used in each direction, the bi-linear approach constructs new data points from a discrete set of

7 original data values based on a quadratic function (Press et al., 1993). The resolution of this

8 approach is further examined in the sections below.

10 3. Results

11 3.1. Maps of Reflection Amplitudes

12 Fig. 2 shows the region of interest in this study. Approximately 5000 high-quality traces are

13 retained after the data selection procedure detailed in the previous section. The ray theoretical

14 reflection points of the precursors (see Fig. 2) provide adequate resolution for the entire study

15 area. Furthermore, the increased data coverage in the latitude range of 35°-50° facilitates a

16 direct comparison of the mantle reflectivity structures in the vicinity of southern/central Japan

17 (cross-sections A and B) with those beneath the Kuril trench (cross-section C).

19 The Amplitude variations of 3D depth-converted SdS waves indicate the presence of large-scale

20 structures in the MTZ and shallow lower mantle. The top of the MTZ (Fig. 3) contains an

21 elongated, highly reflective zone (HRZ), extending from the northern Great Khingan Range in

22 the east to the northwestern corner of the study region beneath the Gobi desert. This 1500-km

23 wide anomaly reaches its maximum amplitude (9% of that of SS, for short, 9%) at ~425-km

8 1 depth, which is approximately 15 km below the global average of the 410-km seismic

2 discontinuity (Fig. 3A) (Gu et al., 2003; Houser et al., 2008). A second, weaker HRZs is visible

3 east of the Wadati-Benioff zone along the Kurile and Japan arcs, peaking at ~8% amplitude near

4 the Hokkaido corner (see Fig. 3A).

6 The HRZs at the top of MTZ decays quickly with depth and the reflectivity pattern at ~520 km

7 depth is dominated by a strong (5-8%), uniquely shaped reflector (Fig. 3B). The center of this

8 reflector is located near Sikhote-Alin Mountains, roughly coinciding with the slab corner

9 between Japan and Kuril subduction zones outlined by Sam Gudmundsson and Sambridge

10 (1998) west of the Hokkaido corner (see depth map at 540 km, Fig. 3B). The orientation of this

11 boomerang-shaped structure (see map at 520 km) changes from ~30 deg oblique to the trench-

12 perpendicular direction west of Honshu Island to trench-perpendicular beneath northeastern

13 China. The vertical dimension of this mid-MTZ HRZ is no greater than 40 km (see Fig. 3B).

15 Large-scale reflective structures are clearly visible at the base of the upper mantle (Fig. 3C) and

16 below (Fig. 3d). Major north-south oriented HRZs are observed at 675-km depth northwest of

17 the Japan-Kuril arc-arc interaction region and the eastern section of the Gobi desert, respectively

18 (see Fig. 3C). The maximum amplitudes of both anomalies exceed 10%. The depths of the

19 HRZs indicate local depressions of 20+ km on the 660 beneath northeastern China. The

20 geographical locations of these HRZs roughly overlap with those of two lower-mantle reflectors

21 detectable at 900-930 km depths. The stronger and slightly deeper of the two HRZs (see 6%

22 amplitude isosurface, Fig. 3D) lies beneath the slab corner between Japan and Kuril subduction

9 1 zones. This semi-linear reflective structure is approximately trench-perpendicular and spans the

2 entire Wadati-Benioff zone in this arc-arc interaction region.

4 3.2. Correlation between reflectivity and seismic velocity

5 Detailed information on the temperature-dependent seismic velocity and impedance-driven

6 reflectivity structure is necessary to accurately characterize mantle structure and processes near

7 subduction zones. To explore wave amplitude vs. velocity relationship, we overlay reflectivity

8 depth cross-sections (Fig. 4; see Fig. 2 for reference) with high-resolution regional P velocities

9 reported by Obayashi et al. (2006). While the use of a regional S velocity model would be ideal,

10 key mantle heterogeneities in the study region are better resolved by the high-resolution P wave

11 tomography (see review by Fukao et al., 2009). Reflections within the depth ranges 120-150 km,

12 380-440 km and 630-700 km are consistently observed in all cross-sections despite substantial

13 lateral variations in depth and amplitude. The focus of this study is on the MTZ and lower

14 mantle where waveform complexities associated with SS sidelobes are minimal (e.g., Shearer,

15 1993; Gu et al., 2003).

17 Fig. 4A shows highly undulating MTZ boundaries between the Pacific Plate and the volcanic arc

18 near central Honshu Island. The 410 east of the Japan trench undergoes 15-20 km local

19 depression relative to the cross-sectional average depth of 415 km.?? This 500-km wide HRZ

20 reaches the maximum reflection amplitude of ~8% beneath central Honshu Island, approximately

21 overlapping with a P wave low-velocity zone centered between 380-400 km depths (Obayashi et

22 al., 2006; see also Zhao and Ohtani, 2009; Li and van der Hilst, 2010; Bagly et al., 2009). The

23 reflectivity structure changes sharply toward the Wadati-Benioff zone where the 410 reaches

10 1 local minima in both depth (~395 km) and reflection amplitude (~5%) (see Fig. 4, Profile A).

2 Complex reflective structures are also evident at the base of the MTZ east of the Japan trench.

3 The 660 shows 25+ km peak-to-peak topography and the undulations appear to negatively

4 correlate with those of the 410 along the trench dip. Major depressions are identified beneath

5 eastern Sea of Japan (~680 km) and Gulf of Chihii (~673 km) (see Figs. 3C and 4, Profile A),

6 with the former showing a slight offset from the center of predicted MTZ high velocities.

8 The shape of the high-velocity structure becomes quasi-linear near northern Honshu Island

9 where a significant number of deep-focus earthquakes have been recorded (Fig. 4, Profile B).

10 The 410 remains depressed in the east of the Wadati-Benioff zone (see profile A). A strong HRZ

11 is visible at ~300 km depth in this region, approximately outlining with the top of the low-

12 velocity zone (also see Fig. 3A) above the MTZ. The reflection characteristics of the 410 are

13 generally consistent with those from profile A, but the lateral variations in amplitude and depth

14 are visibly diminished relative to the former profile. At the base of the MTZ, the 660 shows

15 extreme local topography in the vicinity of the Wadati-Benioff zone. The depth of the 660

16 beneath the island arcs is ~645 km, the shallowest level in the entire profile, which significantly

17 reduces the MTZ width (~225 km) along the trench dip (see Fig. 4, Profile B). This anomalous

18 topographic structure on the 660 is accompanied by a broad depression beneath the Sea of Japan

19 and Changbai hotspot. The 1000-km wide structure west of the Hokkaido corner overlaps with a

20 P wave high-velocity zone near the base of the MTZ, but its lateral dimension is considerably

21 greater than that inferred from the 1+% P velocity variations.

11 1 The high-velocity structure beneath the Kuril subduction zone (Profile C) is visibly more

2 complex than those beneath the Japan subduction system, providing convincing evidence for 1) a

3 fast zone along slab dip that extends down to 750+ km depths, and 2) a horizontal MTZ anomaly

4 west of the Sea of Okhotsk with a possible ‘necking’ beneath the Sikhote-Alin Mountains. The

5 reflectivity structure in Profile C accentuates the complex slab morphology and kinematics in

6 this region. Apparent reflection gaps are observed on the prodominantly continuous 410 and 660

7 along the Wadati-Benioff zone, with the latter anomaly nearly spanning the entire Sea of

8 Okhotsk. The shape of the 660 phase boundary west of this low-amplitude region closely

9 matches the outline of the 1% high-velocity structure in the MTZ (see Fig. 4, Profile C). We also

10 identify a highly undulating, piece-wise continuous lower mantle reflector beneath this profile,

11 showing the largest amplitude (~10%) beneath the reflection gap on the 660. The presence of

12 this lower-mantle reflector and isolated MTZ HRZs will be discussed in detail in Section 4.

14 The cross-sections shown by Fig. 4 (Profiles A-C) paint markedly different pictures of MTZ

15 reflectivity structures between Japan and Kuril subduction zones. A north-south transect over the

16 deepest part of the Wadati-Benioff zones (Fig. 4, Profile D) highlight the key observations that

17 differentiate between these two subduction systems. South of Hokkaido corner, large-scale

18 high-velocity structures appear to reside within the MTZ. Despite slightly reduced amplitudes,

19 the MTZ phase boundaries are generally detected and laterally continuous. In particular, the 660

20 is generally deeper than regional averages and the largest ‘visible’ depressions is detected

21 between the Korea Strait and Sea of Japan. On the other hand, the Kuril subduction zone

22 embodies a vertically continuous high-velocity structure that extends into the shallow lower

23 mantle. This P velocity anomaly is supported by a strong HRZ at ~930 km depth. Furthermore,

12 1 the amplitudes of the MTZ phase boundaries in the same regions are clearly below the threshold

2 of detection using SS precursors. It is worth noting that 1+% P velocities appear to reach a

3 depth of ~280 km, which coincides with a strong, possibly deformed, shallow mantle reflector

4 between the two subduction zones.

6 A common observation between the Japan and Kuril subduction zone is the presence of mid

7 MTZ reflector(s) (see Fig. 4, Profiles A-D). We identify a single HRZ with maximum

8 amplitudes in excess of 6% at ~525 km near or within the Benioff zones in the southern profiles.

9 Two isolated mid-MTZ reflectors are present under the Kuril subduction zone at approximate

10 depth ranges of 500-530 km and 580-600 km, respectively. The depths of these reflectors vary

11 considerably in each profile, whereas the amplitudes generally increase from South to North.

13 3.3. Hypothesis testing and nominal resolution

14 Several procedures are implemented to ensure the stability and accuracy of the migration method

15 as well as the resolution of the SS precursor data set. To investigate the effect of earthquake

16 source and the migration algorithm, we compute synthetic seismograms (Fuchs and Muller,

17 1971; Kind, 1978; Hermann and Wang, 1985) for all source-station pairs based on PREM

18 (Dziewonski and Anderson, 1981) and earthquake source information from GCMT (see also

19 Section 2). The synthetic data set is then subjected to the same filtering, binning and migration

20 procedures as the actual observations. Fig. 5 shows the sample output for Profile C, which

21 validates at least two key premises of this study. First, the two bounding MTZ reflectors are

22 migrated to 400 and 670 km, respectively, to at least two decimal places. These values are

13 1 consistent with those of PREM, the 1D model used in the migration procedure, which suggest

2 that the time-to-depth mapping of the actual data is precise in the absence of lateral variations in

3 velocity or phase boundary topography. Furthermore, the amplitudes and depths of the MTZ

4 phase boundaries are nearly constant along the profile, which imply that the collective influence

5 of earthquake source mechanism, station response, and phase equalization on the results of data

6 migration stacks is negligible along this (see Figs. 5) and other (not shown) profiles.

8 Questions have surfaced in recent years regarding the accuracy of the structure/topography

9 inferred from SS precursors due to the mini-max nature of reflected waves and their wide Fresnel

10 zones at long periods (Neele et al., 1997; Chaljub and Tarantola, 1997). Shearer et al. (1999)

11 addressed some of the potential biases through a multi-scale resolution analysis. By inverting for

12 synthetic differential travel times, they showed that a topographic inversion using long-period SS

13 precursor observations is virtually immune to smaller-scale artifacts at a major subduction zone.

14 Recent high-resolution images from the investigations of subduction slabs (Schmerr and

15 Garnero, 2007; Heit et al., 2010), hot mantle plumes (Schmrr and Garnero, 2006; Gu et al., 2009;

16 Cao et al., 2011) and lithosphere (Rychert and Shearer, 2009) are further testimonies of the

17 strong resolvability of SS precursors on finer-than-expected structures at mantle depths. Shear

18 waves has been known to resolve structures with length scales beyond their ‘nominal’ resolution,

19 especially when waveform information is incorporated (Ji and Nataf, 1998; Mégnin and

20 Romanowicz, 2000). In the case of SS precursors, minor errors are expected when relatively

21 large Fresnel zones of SS precursors collapse onto the fine grid adopted by this study, though the

22 lateral depth/amplitude differences between the averaging centers could persist and the apparent

14 1 connections between reflection amplitude, seismic velocity and seismicity (see Sections 3.1 and

2 3.2) are hard to dismiss as random occurrences.

4 Without repeating the successful experiment performed by Shearer et al. (1999), we examine

5 different CMP sizes to determine the optimal level of tradeoff between stability and resolution.

6 Fig. 6 shows the a comparison of reflectivity maps at 680 km based on averaging bins sizes of 2

7 x 6 = 12 deg2 (Fig. 6A) and 5 x 10 = 50 deg2 (Fig. 6B). Differences in the suggested spatial

8 scales of the anomalies are apparent. A significant number of reflectors, some poorly resolved

9 due to insufficient data, exist in the former map whereas larger bin sizes tends to over-damp the

10 lateral variations in 660 topography. However, the location and maximum amplitudes of major

11 HRZ, e.g., a semi-linear structure across northern Honshu Island and a large, uniquely shaped

12 zone contouring the deepest part of the arc-arc interaction region, are minimally affected by bin

13 sizes. Our final choice of averaging area (32 deg2) represents an effective, albeit subjective,

14 compromise between image stability and resolution.

16 3.4. Uncertainty of reflectivity structure

17 We estimate the uncertainty of the reflectivity profiles based on bootstrapping resampling

18 algorithm (Efron, 1977). For each averaging bin, we first construct a ‘bootstrapped’ data set of

19 equivalent size to the original data set through random drawing. This procedure is performed

20 with the aid of a random generator (Press et al., 1992) and allows for repeated selections of the

21 same seismogram. We then perform data stacking and migration on this simulated data set and

22 obtain a single summary migrated seismogram for this particular averaging bin. This random

23 drawing and migration/averaging procedure is repeated 300 times in the same data gather to 15 1 obtain a statistically significant distribution of reflectivity at each depth. We estimate the

2 effective uncertainty by the standard deviation of these 300 bootstrapped seismograms (Efron,

3 1977; see also Shearer, 1993; Deuss and Woodhouse, 2002; Gu et al., 2003; Lawrence and

4 Shearer, 2006; An et al., 2007; Zheng et al., 2007), and apply the same treatment to all averaging

5 bins along each profile.

7 The bootstrapped reflectivity profiles, which are constructed based on the average of the re-

8 sampled seismograms at each data gather, are nearly identical to the respective profiles shown by

9 Fig. 4. The bootstrapped uncertainties based on one standard deviation (Fig. 7) are generally

10 lower than 3% below 150 km. The spatial variation in uncertainty is nearly random, which

11 implies that the main MTZ reflectivity structures are reasonably well resolved in all profiles.

12 However, all four profiles show a 200-500 km wide section of increased uncertainties (reaching

13 ~3% amplitude) that intercepts the seismogenic zone, e.g., beneath the Japan trench in Profiles A

14 and B and Strait of Tartary in Profile D (see Fig. 7). This anomalous zone is partly caused by

15 relatively sparse data coverage (see Fig. 2), though the scattering associated with inclined high-

16 velocity slab structures cannot be ignored. Further discussions of the latter effect will be

17 provided in Section 4.

19 4. Interpretation and discussion

20 Using reflected/scattered waves to illuminate the shape of major thermal and/or compositional anomalies

21 is analogous to ‘motion tracking’ in animation cinematography. In a nutshell, both procedures take

22 advantage of the relationships between reflection/scattering strengths and changes in material

23 properties including density, bulk or shear modulus and, in the case of motion tracking, index of

16 1 refraction of electromagnetic waves. Signals analyzed by both applications are strongly sensitive to

2 gradients in material properties and, with proper treatments, can be used to decipher the shape or

3 movements of the enclosed mass. On the other hand, destructive interference or scattering of the

4 waves caused by structural asperities could present challenges, albeit providing additional

5 information, to both applications. The incorporation of additional physical constraints could be

6 highly beneficial. For the case reflectivity imaging, the combination of reflectivity imaging and

7 seismic tomography can substantially improve our existing knowledge on morphology of subducted

8 crust and lithosphere in the northwestern Pacific region.

10 Many important factors must be considered in the discussion of the morphology and kinematics

11 of subducting slabs. From a mineralogical viewpoint, the slab geometry and the width of the

12 MTZ are strongly influenced by mineralogical phase transformations of olivine to wadsleyite

13 (near 400 km), wadsleyite to ringwoodite (near 520 km), and ringwoodite to perovskite +

14 magnesiowustite (near 660 km) (Katsura and Ito, 1989; Ita and Stixrude, 1992; Helffrich, 2000;

15 Bina, 2003 and references therein; Akaogi et al., 2007). The endothermic phase change at the

16 base of the MTZ increases local buoyancy forces, which can deflect subducting slabs and aid its

17 stagnation within the upper mantle (Christensen, 1995; Billen, 2008, 2010; Fukao et al. 2009).

18 Under thermodynamic equilibrium, a cold, water-rich slab is expected to raise the 410, depress

19 the 660 (due to the opposite signs of their Clapeyron slopes), and be responsible for a wide range

20 of reflective bodies within the mantle. The presence of water can strongly impact the phase

21 changes in the MTZ (e.g. Inoue et al. 1995; Kohlstedt et al., 1996; van der Meijde, 2003; Ohtani

22 et al. 2004; Kombayashi and Omori, 2006; Huang et al., 2006; Litasov et al., 2006; Suetsugu et

17 1 al., 2006). Below is a detailed account of some of the observed reflectivity structures in the

2 general framework of MTZ mineralogy and temperature.

4 4.1. Amplitudes of the MTZ discontinuities

5 The amplitudes of the reflections from the MTZ phase boundaries are functions of the impedance

6 contrast across the reflecting surface and the transition width. Furthermore, due to the

7 summation of multiple seismograms at each location and the use of SS amplitude as the

8 normalization term, the topography on the interface and regional variations of SS can also

9 significantly impact the relative SS precursor amplitudes. This study exclusively focuses on the

10 positive reflections associated with increased material impedances with depth. This is a

11 subjective decision prompted by the simple observation that the signs of well-resolved

12 reflectivity structures are predominantly positive in our study area. Admittedly, many positive

13 phases are accompanied by sizeable negative peaks that could result from reductions in velocity

14 and/or density, e.g., near the top of a low velocity zone or the bottom of a high velocity structure.

15 We defer discussions of negative phases to a future study.

17 The detectable ranges of amplitudes are 4-9% for S410S and 4-12% for S660S, both showing

18 significant lateral variations. The former range overlaps with the predicted values of ~8% from

19 PREM (Dziewonski and Anderson) and global average of 6.7% (Shearer, 1996) based on SS

20 precursor observations, whereas the latter range falls well short of the predicted 14% (Shearer,

21 2000). These individual amplitude estimates are strongly affected by the strength of SS, the

22 normalizing reference phase. For instance, the presence of attenuating low-velocity structures

23 (e.g. Zhao et al., 1992, 1997, 2004; Lei and Zhao, 2005; Huang and Zhao, 2006), especially near

18 1 back arc regions (e.g., Xu and Wiens, 1997; Roth et al., 1999, 2000), could reduce the absolute

2 amplitude of SS and increase the relative amplitude. Compositional variations associated with

3 Al at the base of upper mantle (e.g., Weidner and Wang, 1997, 2000; Deuss and Woodhouse,

4 2002; Deuss, 2009) or Fe content (Akaogi et al., 2007; Inoue et al., 2010) are also known to

5 broaden phase boundary widths and cause reductions in precursor amplitudes. A more stable

6 parameter is the amplitude ratio between the 410 and the 660 (e.g., Shearer, 2000), which we

7 estimate to be within the range of 0.7-0.8. This value is slightly higher than the earlier estimates

8 of 0.64-0.68 based on global SS precursor (shearer, 1996) and regional ScS observations

9 (Revenaugh and Jordan, 1991), but it is in poor agreement with that of PREM (0.5). A regionally

10 sharp 410 (e.g., Benz and Vidale, 1993; Vidale et al., 1995; Neele, 1996; Melbourn and

11 Helmberger, 1998; Ai and Zheng, 2003; Jasbinsek et al., 2010) could , although the presence of a

12 fluid-rich lens near the 410 (Smyth and Frost, 2002; van der Meijde, 2003; Inoue et al., 2010).

13 While these effects are difficult to constrain reliably based on seismic observations, scattering

14 associated with undulations on the two MTZ bounding discontinuities are more readily

15 observable (Shearer, 2000). The presence of dipping structures, particularly in the vicinity of

16 slabs, can preferentially lower the ‘perceived amplitude’ of the 660, hence the amplitude ratio of

17 410 vs. 660, due to the 25-30% larger topography on the 660 relative to that on the 410 (see Figs.

18 3 and 4). The following sections carefully examine discontinuity depths and their implications

19 for slab geometry and dyanmics.

21 4.2 Depth correlation of the MTZ discontinuities

22 The migrated reflectivity profiles provide new insights on the effect of mantle temperatures on

23 phase boundary variations. Results from high-pressure mineral physics (e.g., Katsura and Ito, 19 1 1989; Ita and Stixrude, 1992; Irifune et al., 1998; Helffrich, 2000; Akaogi et al., 2007) have

2 predicted a negative correlation, hence an increased transition width, between the phase

3 boundary undulations in an olivine-dominated mantle. Seismic evidence from regional (e.g., Li

4 et al., 2000; Collier et al., 2001; Lebedev et al., 2002; Saita et al., 2002; Ai et al., 2003; van der

5 Meijde et al., 2005; Ramesh et al, 2005; Tonegawa et al. 2005) and global (Shearer and Masters,

6 1992; Shearer, 1993; Gossler and Kind, 1996; Gu et al., 1998; Flanagan and Shearer, 1998;

7 Lawrence and Shearer, 2006; Houser et al., 2008) analyses have generally supported this

8 hypothesis, but analyses based on lower-resolution approaches have largely attributed the

9 increased thickness to a strongly deformed 660 that correlates with the thermal variations at the

10 base of the upper mantle (Flanagan and Shearer, 1998; Gu et al., 1998, 2003; Gu and

11 Dziewonski, 2002; House et al., 2008). The depth of the 410 remains problematic in view of

12 mantle chemistry on all scales (e.g., Gilbert et al. 2002; Fee & Dueker 2004; Du et al. 2006; Gu

13 and Dziewonski, 2002; Gu et al., 2003; Deuss, 2007; Schmerr and Garnero, 2007; Tauzin et al.,

14 2008). Additional assumptions involving corrections (Flanagan and Shearer, 1998; Gu et al.,

15 2003; Schmerr and Garnero, 2006; Deuss, 2007; Houser et al., 2008) and/or mechanisms

16 predicated on extensive compositional variations (Schmerr and Garnero, 2007; Deuss, 2007; Gu

17 et al., 2009; Houser and Williams, 2010) are needed to reduce the difference between observed

18 and expected MTZ phase boundary perturbations.

20 To examine the correlation between temperature and discontinuity topography in our study area,

21 we focus on Profile A where both the 410 and 660 show the largest detectable topographic

22 variations and amplitudes near the Wadati-Benioffz zone (Fig. 8). The respective peak-to-peak

23 depth variations of the 410 and 660 are approximately 30 km and 410 km, which are comparable

20 1 to the largest variations reported by earlier global studies (Shearer, 1993; Gossler and Kind,

2 1997; Flanagan and Shearer, 1998; Gu et al., 2001, 2003; Houser et al., 2008; Lawrence and

3 Shearer, 2008). Both phase boundaries undergo extreme deformation from the trench onset to

4 the deepest part of the Wadati-Benioff zone across southern Japan (see Fig. 8A). A simple bin-

5 by-bin correlation assuming vertical thermal structures, the same approach used in the

6 aforementioned global studies, suggests a positive correlation between discontinuity depths over

7 the length of the profile (see Fig. 4). To account for non-vertical structures following the slab dip

8 (~30 deg, Gudmundsson and Sambridge, 1998), we revise the correlation analysis by applying an

9 indexing change such that the depth of the 410 at a given location is correlated with the 660

10 depth at a location ~200 km further inland. The dip-corrected phase boundaries show clear

11 negative correlation in the vicinity of the slab (see Fig. 8A) and the corrected correlation

12 coefficient is -0.4 for the entire profile, a statistically significant value that clearly favors a

13 thermal origin for the observed MTZ topography. A key reason for the strong negative

14 correlation is the observed elevation of the 410 within the Wadati-Benioff zone. This feature

15 represents a major departure from those of earlier time-domain global studies of SS precursors

16 (e.g., Flanagan and Shearer, 1998; Gu et al. 2003), which we attribute to improved data

17 resolution in this study. From a broader perspective, this experiment not only highlights the

18 ability of SS precursors in resolving small-scale subduction zone anomalies, but also provides a

19 blueprint for to improve global correlation analyses via a priori information such as slab dip

20 angles.

22 4.3. Continuity of the 410 beneath northeast China

21 1 There have considerable discussion of results obtained from laboratory experiments on the existence and

2 support for a water/melt rich layer near the top of the MTZ (Wood, 1995; Inoue et al., 1995, 2010;

3 Kohlstedt et al., 1995; Smyth and Frost, 2002; Frost and Dolejs, 2007). Based on these studies,

4 wadsleyite has a strong capacity to accommodate hydroxyl (OH−), storing up to 3 wt.% H2O under

5 equilibrium conditions (Wood, 1995; Inoue et al., 1995, 2010; Smyth and Dolejs, 2007). These

6 laboratory-based measurements have been supported by regional (e.g., Revenaugh and Sipkin, 1994;

7 Zheng et al., 2007; Schmerr and Garnero, 2007; Schaeffer and Bostock, 2010) and global (Tauzin et

8 al., 2010) seismic observations of low-velocity zones at similar depths that cannot be sufficiently

9 explained by thermal variations. The infiltration of hydrous melt is further constrained through

10 geodynamical calculation and synthesis (e.g., Bercovici and Karato, 2003; Karato, 2006; Leahy and

11 Bercovici, 2007, 2010).

13 Our migrated reflectivity structures provide further regional constraints on this hypothesized hydrous

14 layer above the 410. The 410 west of the Wadati-Benioff zone (Fig. 8B) is consistently shallower

15 than the regional average in this study. The largest topography is observed in the southernmost

16 cross-section, reaching a depth of ~400 km beneath Korea and northeastern China. The two northern

17 profiles B and C show modest highs of ~410 km in the topography of the 410 near the Changbai

18 hotspot and Sikhote-Alin Mountains, respectively. The average amplitudes of the 410 in all three

19 profiles far exceed the regional average, despite visible falloffs in the middle of the highlighted

20 section in the latter two profiles (see Fig. 8B). These characteristics are reminiscent of those

21 reported beneath the Tonga subduction zone (Zheng et al., 2007) based on migrations of precursors

22 to both P and S depth phases. However our highlighted section shows strong positive reflections,

23 which is opposite to those reported near Tonga, and the perturbations in depth (<15 km relative to

24 410 km) is weaker than those presented by the earlier study (>20 km). Our highlighted region (see

25 Fig. 8B) is also farther away from the Wadati-Benioff zone than the target area in Zheng et al.

22 1 (2007), though metasometism involving slab-derived fluids rising through the flattened part of slabs

2 (see Fukao et al., 2009 for review) could potentially be as extensive as that beneath slab wedge. In

3 fact, intraplate volcanoes nears Changbai mountains and Wudalianchi region (see also Fig. 8B,

4 Profile B) have been closely linked to processes similar to back-arc spreading of the Japan slab (Lei

5 and Zhao, 2005; Huang and Zhao, 2006).

7 Schmerr and Garnero (2007) present another intriguing comparison. Based on multiple cross-sections in

8 South America, this earlier study inferred a ‘melt lens’ based on evidence of delayed and

9 split/missing S410S reflections east of the Nasca-South America convergent zone. The presence of

10 highly anomalous underside reflections received further support from Contenti et al. (submitted,

11 2011) based on the method presented in this study. However, the complexity of the S410S signal

12 from South America far exceeds that from northeastern China. Should a fluid-rich layer be present

13 atop the MTZ beneath our study region, its spatial scale, infiltration/storage mechanism and/or

14 chemistry are likely to be different from those near Tonga and South America subduction systems.

16 4.4. Slab stagnation and distortion

17 Subducted ocean basins in the western Pacific region have been known to deflect to a near-horizontal

18 direction the MTZ for nearly two decades (Okino et al., 1989; van der Hilst et al., 1991; Fukao et al.,

19 1992, 1993). Since then, ample evidence of slab stagnation (Fukao et al., 1993, 2001) in subduction

20 zones worldwide has been provided by global and regional tomographic images with improved

21 accuracy and resolution (Fukao et al., 2001, 2009; Zhao and Ohtani, 2009; Li and van der Hilst, 2010;

22 Sugioka et al., 2010) and anomalous dip-angle variations suggested by the distribution of

23 intermediate-depth earthquakes (Chen et al., 2004). The conditions and characteristics of stagnant

24 lithosphere have been constrained further through numerical calculations incorporating thermo-

23 1 petrological buoyancy forces (Tetzlaff and Schmeling, 2000; Bina et al., 2001; Bina and Kawakatsu, 2010), rheology (Billen and

2 Hirth, 2007; Billen, 2008), and plate history and rollback (Torii and Yoshioka, 2007; Christensen, 2010; Zhu et al.,

3 2010).

5 6 With the help of seismic velocities, the reflectivity information provided by our study can place

7 crucial constraints on slab deformation at the base of the MTZ and the shallow lower mantle. In

8 particular, the shape of the HRZs near the 660 provides useful measures for the geometry and

9 dimension of the stagnant slabs. The two southern profiles presented in Fig. 3C and Fig. 4A-C

10 consistently show two distinct zones of large-lateral scale depression (Fig. 9), 1) near the

11 piercing point of the slab at the base of upper mantle, and 2) in the second half of the stagnant

12 slab inferred from recent tomographic models (e.g., Huang and Zhao, 2006; Fukao et al., 2009).

13 The two depressive zones have nearly identical shapes, particularly in Profile B, and depth of the

14 660 between them ranges from 655 to 660 km in both cases. Profile A shows significantly larger

15 topography than Profile B near the slab piercing point. For an isochemical mantle, the maximum

16 depth of ~685 km would suggest a temperature increase of XX-XX deg C depending on the

17 selected Clapeyron slope (REF). The reduction in topography from south (Profile A) to north

18 (Profile B) along the island arcs is in general agreement recent studies based on receiver

19 functions (Niu et al., 2005) and postcursors to sScS (Yamada and Zhao, 2007). The reduced

20 horizontal gradient in the topography of the 660 beneath northern Honshu could be caused by a

21 ‘soft’ slab (Li et al., 2008) under the influence of trench migration and rollback. However, Li et

22 al. (2008) detected little or no oceanward broadening of the 660 from high-resolution S to P

23 converted waves. This is inconsistent with the apparent shift between the high-velocity contours

24 and the onset of the depressive zones in the vicinity of the island arcs (see Fig. 4 and Fig. 10). 24 1 Resolution differences of the two data sets (SS precursors vs. receiver functions) may be a

2 contributing factor, still, 100-300 km horizontal broadening/ponding of the Pacific slab at the

3 base of MTZ in the oceanward direction remains a strong possibility.

5 A dimensional analysis of slab geometry based on the topography on the 660 is informative but

6 requires subjective definitions. Assume the points of intersection at 670 km depth mark the

7 corners of the topographic structures, we estimate the horizontal dimensions of depressive zones

8 to be 350-450 km in Profile A and 500-600 km in Profile B. The respective topographic highs

9 between the depressions are estimated to be ~700 km and ~400 km. The total length beyond the

10 depressions near the slab piercing point is approximately 1050 km for Profiles A and 900 km for

11 Profile B. These values are reasonably consistent with the estimated length of 800-1000 km for

12 deflected slab bodies (Huang and Zhao, 2006; Fukao et al., 2010), especially if slight reductions

13 due to horizontal averaging are considered in our estimates. However, as suggested by Fig. 10

14 and the estimates above, the truly ‘flat’ part of the slab that depresses the 660 phase boundary is

15 most-likely less than 600 km in width.

17 The migration-based topography of the 660 (see Fig. 9) challenges the ‘flatness’ of stagnant slabs. The

18 observation of contention is the average or shallow 660 between the depressive zones, particularly in

19 Profile A, whereas broad, continuous depression zones have been reported earlier though seismic

20 tomography (see Fukao et al., 2009 for review) and reflection depth/MTZ thickness imaging (e.g.,

21 Shearer and Masters, 1992; Flanagan and Shearer, 1998; Gu et al., 1998, 2003; Lawrence and

22 Shearer, 2006; Houser et al., 2008). Furthermore, the amplitude of the 660 within this uplifted

23 region is consistently higher than the regional averages, which is consistent with the expected

25 1 decrease of ringwoodite-perovskite+magnisiowustite phase loop under high-than-average

2 temperatures. The observed phase boundary behavior is plausible based on recent geodynamical

3 calculations of slab geometry that consider 1) trench retreat (Christensen 1996; Tagawa et al. 2007; Zhu et al.,

4 2010) or 2) temperature- and pressure-dependent viscosity (Karato and Wu 1993; see Fig. 12 of Fukao et al. 2010) .

5 These calculations infer distinct zones of depression at the slab piercing and re-entry points, between

6 which the 660 remains largely unperturbed. The images provided by these models are consistent

7 with our observations in the MTZ, though the expected reflections from the horizontally oriented

8 slab segment in the shallow lower mantle (e.g., Fukao et al., 2009) are not clearly observed from our

9 data set (see Fig. 9).

11 Alternatively, the internal undulations within stagnated slab body could suggest vertical deformation of

12 slab interface in the MTZ. Part of the lateral variations may be related to advection (Kellogg et al.

13 1999; Obayashi et al., 2006), where the ambient and relatively hot mantle material got ‘trapped’

14 during the interaction between the tip of the downgoing slab and viscous lower mantle. Trench

15 migration and rollback history could play a major role, as the current geometry of stagnant slab could

16 reflect changes in slab dip over the course of 100+ Ma (see Schmid et al., 2002 for the case of

17 Farallon plate subduction). Finally, the presence of water (e.g., Listov et al. 2002, 2006; Inuoe et al.,

18 2010) and possible separation of oceanic crust from the downgoing lithosphere (Irifune and

19 Ringwood, 1995;van Keken et al., 1996; Hirose et al., 1999, 2005) could also contribute to strong gradients in the

20 topography of the 660 within the ‘flat’ part of the slab.

23 4.5. Slab penetration beneath Kuril subduction zone

24 The reflectivity structures add new insights into the long-standing debate about the depth of slab

25 in the Pacific northwest (van der Hilst et al., 1991; Fukao et al., 1992; van der Hilst et al., 1997;

26 1 Fukao et al., 2001, 2009). While the vertical extent of slabs and the general style of mantle

2 convection remain debated on the global scale, there is growing evidence of scattered and

3 deformed slab material in the lower mantle (van der Hilst et al., 1997; Bijwaard et al., 1998;

4 Fukao et al., 2001, 2009; Obayashi et al., 2006; Courtier and Revenaugh, 2008; Li and van der

5 Hilst, 2010; Chang et al., 2010).

7 Among the various HRZs documented in this study, MTZ anomalies contained in Profiles C and D

8 provide strong evidence for penetrating slabs in the western Pacific region. The most visible change

9 in the reflectivity structures from central Honshu slab to southern Kuril slab is the amplitude

10 reduction of the 410 and 660, highlighted by the apparent reflection gaps in Profiles C and D. These

11 gaps coincide with the Wadati-Benioff zone of the Kuril slab and their lateral dimensions reflect the

12 increasing width of the high velocity structure from the top to the bottom of the MTZ (see Fig. 10A).

13 The origin(s) of these reflection gaps remain(s) debatable. Factors that have considerable impact on

14 the amplitudes of the MTZ reflectors (see also Section 4.1) include Al, water and Fe contents and

15 optics.

17 There are merits and significant caveats in attributing the observed reflection gaps to variations in mantle

18 chemistry (e.g., the first three factors listed above). Under proper mantle conditions, an increase in

19 Al content could broaden the depth range of garnet-to-perovskite transformation and influence

20 olivine and pyroxene normaltive proportions near the base of the upper mantle (Gasparik, 1996;

21 Weidner and Wang, 1998; 2000). In a low temperature regime, e.g., subduction zones examined in

22 this study, majorite garnet (a Al bearing mineral group) can transform to metastable ilminite that

23 eventually transforms to Ca-perovskite (e.g., Weidner and Wang, 1998). These phase transitions

24 exhibit different phase boundary behaviors from the olivine system and adversely impact the

27 1 interpretation of discontinuity depths and amplitudes. The presence of Al-bearing Akimotoite could

2 introduce further complexities, e.g., a high velocity layer or a steep velocity gradient, to mid MTZ

3 depths at low temperatures (Gasparik, 1996; Wang et al., 2004). However, changes in Al content

4 mainly impact mantle reflectivity structure under mid-to-lower MTZ pressure-temperature

5 conditions (e.g., Weidner and Wang, 2000; Wang et al., 2004). The restrictive condition greatly

6 weakens the role of Al in view of the unexplained absence of the 410 within Kuril slab.

8 Water transported into the MTZ by the subducting slab could also modify the impedance contrast, hence

9 the visibility of a reflecting body (van der Meijde, 2003; Ichiki et al., 2006). Aided by strong

10 capacities of wadleyite and ringwoodite to retain water (Inoue et al. 1995, Kohlstedt et al. 1996; see Fukao et al., 2009

11 for review), a hydrous MTZ can simultaneously affect the width and depth of the 660 (Litasov et al.,

12 2006; Akaogi et al., 2007; Inoue et al., 2010). However, the effect of water on the phase phase loop

13 of the olivine-Wadsleyite transition is rather complex and relatively minor with1 wt% H2O (Inoue et

14 al., 2010). The implication is that a large amount of water must be present in the descending slab to

15 diminish the amplitude of S410S below the detection threshold. Unfortunately, recent seismic

16 observations (Fukao et al., 2009; Bina and Kawakatsu, 2010), particularly those based on a novel

17 modeling strategy for MTZ water content (Suetsugu et al., 2006, 2010), have largely inferred ‘dry’

18 (e.g., <0.5%, Suetsugu et al., 2010) slabs in various parts of the Pacific rim. Mechanism(s)

19 predicated on increased Fe content in slabs are similarly flawed. While increasing the Fe number

20 can substantially broaden the phase loops of both olivine-wadsleyite and ringwoodite-

21 perovskite+magnisiowustite transitions (Litasov et al., 2006; Akaogi et al. 2007; Inoue et al., 2010),

22 the observational support for the enrichment of Fe in subduction zones is not well established.

24 The observed reflectivity gaps are best explained by effects commonly observed in optics. Similar to the

25 scattering of light, the observed amplitudes of the underside SH-wave reflections are strongly

28 1 influenced by the geometry of the reflecting surface. A dipping structure or interface generally

2 causes defocusing or scattering that, depending on the size of the structure relative to the wavelength

3 of the incoming wave, can result in the destructive interference of the reflected/scattered waves.

4 Therefore, local topography on the two MTZ bounding phase boundaries in response to thermal

5 and/or compositional variations are expected to tradeoff with reflection amplitude obtained through

6 averaging. This effect was documented by Chaljub and Tarantola (1997) based on results from

7 finite-difference modeling of S660S amplitude in response to local topography and higher-than-

8 average velocities, though the conclusions of that study has been a subject of considerable debate

9 (e.g., Shearer et al., 1999). We hereby quantify the relationship between topography and SS

10 precursor amplitude based on simulations of stacked SS precursors from a depressed zone assuming

11 uniform (case 1, Fig. 10A) and more extreme (case 2, Fig. 10A) spatial distributions of reflection

12 points. Reflectivity synthetic seismograms (Randall are computed for common explosive source

13 recorded by a station at 130-deg epicentral distance. This experiment is repeated for depth

14 perturbations (positive for the 410 and negative for the 660) ranging from 0 (unperturbed PREM

15 model) to 40 km. The resulting stacked waveforms of SS precursors show a steady decay with

16 increasing vertical topography, particularly for case 2 where the reflection-point distribution is sparse

17 (Fig. 10A and 10B). For both cases, the amplitude drops to 50% for undulations of 15-25 km on the

18 410 and 25-35 km on the 660, which will be problematic during the detection of large topographic

19 features. Between the two phase boundaries, the influence of btopography is larger for the 410 than

20 the 660 due to a smaller assumed velocity jump at the former interface (see Fig. 10B). The

21 amplitude decay could be more severe for sparsely populated data (see case 2 simulations, Fig. 10).

22 Furthermore, the presence of large topography can significantly modify the waveform characteristics

23 of the superimposed seismogram. The wave shape broadens within increasing topography and,

24 depending on the frequency, can split into separate low-amplitude arrivals reflecting the top and

25 bottom of the topographic structure, respectively (see Fig. 10). 29 1

2 An underpinning message from Fig. 10 is that the maximum depth of the 660 could be 700 km or deeper

3 in the Pacific northwest (e.g., Revenaugh and Jordon, 1989; Niu et al., 2005). Based on the

4 impedance contrasts suggested by PREM, the amplitudes of both phase boundaries could easily fall

5 below the detection threshold of ~4% during the migration procedure when the topography exceeds

6 35 km for the 660 and 20 km for the 410. While this is the ‘worst case’ scenario that assumes the

7 averaging bin size is equivalent to the surface area of the topographic structure, it does provide a

8 viable explanation for the missing 410 and 660 within the Kuril slab. The waveform splitting

9 phenomenon (see Fig. 10) also has significant implications for the detection of double reflectors.

10 For example, results from high-pressure mineral physics (e.g., vacher et al., 1998; Weidner and

11 Wang, 1998, 2000; Akaogi et al., 2002) have provided solid laboratory evidence for garnet-ilmenite-

12 perovskite transition near the base of the upper mantle. Within low-temperature slabs, these garnet-

13 related transitions are expected to take place over 60-100 km range in depth (Vacher et al., 1998;

14 Akaogi et al., 2002) that are capable of generating mild reflections in seismic waves.

15 Observationally, the occurrences multiple reflectors have been reported under different tectonic

16 settings (e.g., Deuss and Woodhouse, 2002; Ai and Zheng, 2003; Tibi et al., 2007), but their presence

17 beneath northwest Pacific have been questioned (Lebedev et al., 2002; Tonegawa et al., 2005; Niu et

18 al., 2005). In this study, only Kuril slab (Profiles C and D, Fig. 4) show strongly dipped, weak

19 reflecting bodies centered at ~700- and 780-km depths along the slab dip. These minor reflectivity

20 structures are barely detectable, showing ~4% amplitude each. While it is tempting to link these

21 secondary structures to multiple phase transitions, our numerical experiment above also cautions that

22 the waveform complexities associate with steep topographic structures should be considered in the

23 interpretations.

25 The presence of a high-amplitude lower mantle HRZ beneath Kuril slab (Fig. 11) provide potentially

30 1 crucial support for the vertical extension of Kuril slab beyond the 660. Phase transitions of Ca-

2 perovskite (Stixrude et al., 2007), metastable garnet (Kawakatsu and Niu, 1994; Kubo et al., 2002),

3 as well as transformations of dense hydrous magnesium silicates under lower-mantle pressure-

4 temperature conditions, have been suggested as the origins of a series lower-mantle reflectors (Shieh

5 et al., 1998; Ohtani, 2005; Richard et al., 2006 and references therein). The association of lower-

6 mantle reflectors with phase changes is partially supported by the local maxima of reflection

7 amplitude beneath the reflection gap on the 660. However, reflections from a sub-horizontal lower-

8 mantle HRZ in northeast China between 850-1000 km depths present a potential counter argument.

9 The existence of a chemical boundary (Wen and Anderson, 1997), which would influence the

10 convective flow of mantle, cannot be ruled out.

12 We interpret the presence of the lower mantle reflector as an integral part of an ‘avalanching’

13 slab (Tackley, 1993) based on the following observations: 1) slab gaps at the 410 and 660 that

14 imply substantial mass and heat flux, 2) correlated fast velocity structure that maintains a strong

15 amplitude to depths comparable to that of the lower mantle reflector, 3) the presence of a strong

16 (if not the strongest) lower mantle reflector in the vicinity of the slab gap. These observations

17 are self-consistent and could result from the same process (i.e., slab penetration) under different

18 pressure-temperature conditions and, possibly, mantle chemistry. Since the lower mantle HRZ

19 resides directly below the 660 reflection gap (rather than along the slab dip), the responsible

20 velocity/density structure could have undergone retrograde motion during its descend into the

21 lower mantle. These observations collectively defines the large difference between Kuril and

22 Honshu slabs in terms of maximum vertical extension.

31 1 4.6. Other HRZs and potential inferences

2 Two additional anomalous reflectivity structures from the SS migration images could have significant

3 implications for the mantle structure, dynamics and/or mineralogy if confirmed. First, we identify

4 one (Japan subduction zone) or multiple (Kuril region) mid-MTZ HRZ(s) with reflection amplitudes

5 of 5-9% within the MTZ (see Fig. 3B and Fig. 4). With an exception of one instance east from the

6 slab (see Fig. 4, Profile A), these HRZs are consistently detected within the slab contours suggested

7 by Obayashi et al. (2006). Reflective structures near 520-km depth have been documented nearly 3

8 decades ago in the Pacific northwest from travel time observations (Fukao et al., 1977). It was later

9 proposed to be a mild global seismic discontinuity based on pioneering studies of SS precursors

10 (Shearer, 1990, 1991). Bock (1994) explained this reflector as a potential data processing artifact

11 due to strong low-frequency side-lobes of S410S and S660S phases, though more recent results

12 based on reflected and converted body waves (Gossler and Kind, 1996; Shearer, 1996; Flanagan and

13 Shearer, 1998; Gu et al. 1998, Chevrot et al., 1999; Deuss and Woodhouse, 2001; Gu et al., 2003;

14 Lawrence and Shearer, 2006, Deuss, 2009) have favored an explanation that involves regionally

15 variable, highly undulating reflective structure(s) in the MTZ. In terms of mineral physics, this

16 interface has been attributed to wadsleyite to ringwoodite (Helffrich, 2000, Bina, 2003) and/or

17 garnet to Ca-perovskite (Ita and Stixrude, 1992) phase transitions. In cold mantle regions such as

18 subduction zones, these transformations likely occur at different MTZ depths (Saikia et al., 2008)

19 and produce multiple reflectors (Deuss and Woodhouse, 2001; Deuss, 2009). This may be the case

20 for the observed HRZs within the Kuril slab. Alternatively, delayed meta-stable olivine phase

21 transition (Sung and Burns, 1976; Iidaka and Suetsugu, 1992; Jiang et al., 2008, Bina and Kawakatsu;

22 2010) and the presence of water within slabs are also viable source of enhanced reflections in

23 active plate convergence zones. Ultimately, an accurate interpretation of the anomalous HRZs

24 within the MTZ is predicated upon a greater consensus on the mantle condition surrounding slabs,

32 1 for example, the water content. In view of the apparent north-to-south difference between Japan

2 (a single 520 reflector) and Kuril (multiple reflectors) subduction zones, a combination of these

3 mechanisms may be needed to properly explain our observations in the Pacific Northwest.

5 Lastly, a narrow MTZ and a series of strong HRZs east of the Benioff-zone (see Figs. 4) both

6 suggest low MTZ temperatures. This interpretation is supported by findings in recent studies of

7 ScS reverberations (Revenaugh and Sipkin, 1994; Bagley et al. 2009), seismic tomography

8 (Obayashi et al., 2006; Huang and Zhao, 2006; Zhao and Ohtani, 2009), and electrical

9 conductivity (Ichiki et al., 2006). Furthermore, the strong reflection from these structures (8-

10 12% of SS) may not be sufficiently explained by a thermal origin alone. Compositional

11 variations associated with a hot mantle plume, which was once active during the past 130 Ma,

12 could provide the additional source material necessary to accomodate some of the strong

13 reflections detected in the depth range of 250- 700 km (see Figs. 3 and 4) (Obayashi et al. 2006;

14 Honda et al., 2007; Bagley et al. 2009; Li and van der Hilst, 2010).

17 Conclusions

18 The dynamic processes beneath northwestern Pacific are only a microcosm of those beneath

19 many subduction systems globally. For this reason, inferences based on our high-resolution

20 reflectivity images could be potentially applicable to other regions with similar tectonic settings.

21 Based on the spatial correlation between reflectivity and seismic velocity, we conclude that the

22 origins of the majority of highly reflective zones are thermal, instead of compositional, in nature.

33 1 The combined reflectivity and velocity information enables us to detect and interpret the

2 geometry and strengths of major mantle heterogeneities in the approximate depth range of 300-

3 1000 km.

4 shows clear signs of bending within the MTZ, but the center of the stagnant section of the slab

5 appears to be deformed or folded, as suggested by an average or shallow 660. The depths of the

6 two MTZ bounding olivine phase boundaries are negatively correlated if slab dip is considered.

7 We also identify strong seismic reflector(s) within the slab body within the MTZ through out the

8 The Honshu slab does not appear to extend below the transition zone. negative overall

9 correlation between the depths of the two major olivine phase boundaries. However, localized

10 topography on the 660 within the presumed stagnant part of the slab suggests significant vertical

11 deformation near the base of the upper mantle. A single reflector is identified at the depth range

12 of 500-540 km, which could be associated with by changes in T . which causes strong negative

13 correlations of the olivine phase boundary but and Kuril slabs In particular, our analysis

14 demonstrated that ‘gaps’ in the reflection amplitudes of the 410 and 660 are potentially

15 interconnected with anomalous lower mantle reflectors. Major mass/heat fluxes, large

16 topography on the base of upper mantle, and lower-mantle thermal/composition variations would

17 be expected at these locations. Intermittent reflections within the MTZ offer additional

18 information on the geometries and dynamics of stagnant slabs. In other words, a self-consistent

19 model of mantle processes beneath subduction zones is tenable from the presence, strengths, and

20 depths of mantle reflectors and their spatial correlations with seismic velocities.

22 From a technical standpoint, the results presented in this study provide a glimpse of the future for

23 regional-scale analysis based on intermediate-period SS precursors. Increasingly diverse

34 1 applications in recent years (e.g., Schmerr and Garnero, 2006, 2007; Houser et al., 2008;

2 Lawrence and Shearer, 2008; Gu et al., 2009; Rychert and Shearer, 2009; Heit et al., 2010; Cao

3 et al., 2010; Houser and Williams, 2010) have underlined the remarkable resolving power of this

4 data set, one that was traditionally tapped as a ‘low resolution’ constraint on mantle structure.

5 This trend will likely continue in the foreseeable future, especially in view of the growing

6 number of global seismic networks and applications of array methods.

8 Acknowledgement

9 We sincerely thank Suzan van der Lee for her constructive scientific input to this study. We are grateful

10 to Peter Shearer for his patience and professionalism in handling this manuscript. This study also

11 benefited from the helpful comments and suggestions from Nicholas Schmerr and an anonymous

12 reviewer, as well as from the technical assistance from the IRIS Data Management Center. This

13 project is jointly funded by CFI, Alberta Innovates, Alberta Geological Survey, National Science and

14 Engineering Council (NSERC), and the University of Alberta.

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27 Fig. 1: (A) A schematic drawing of SS precursor reflection from a subducting oceanic lithosphere

28 at the base of upper mantle. These waves are sensitive to the depth and impedance contrast of a

29 mantle interface. (B) Ray theoretical surface reflection points of 6014 high-quality SS waves

30 used in this study. The main tectonic elements and plate boundaries (Bird, 2003) and slab

31 contours (Gudmundsson and Smabridge, 1998) are shown by thick and thin black lines, 51 1 respectively. The surface projections of five mantle transects (see main text) are labeled A-E,

2 extending from central Honshu Island (A) to central Kuril Arc (E).

4 Fig. 2: Key steps in the time-to-depth migration of SS precursors. By placing the aligned SS

5 precursors at the surface (Middle), time samples of transversely polarized seismograms prior to

6 the arrival of SS (Left) can be effectively mapped to corresponding reflection depths (2nd to the

7 Right) along the predicted differential time curves based on PREM (Dziewonski and Anderson,

8 1981). The Right-most panel shows the isotropic shear velocities of PREM down to 1800-km

9 depth.

11 Fig. 3: Interpolated reflectivity maps of SS precursor amplitude variations at MTZ (A to C) and

12 (D) shallow lower mantle depths. An isosurface (threshold = 7.5%) is used to define HRZ in all

13 panels.. The anomalies marked with green dashed lines are discussed in the text. Slab depth

14 contours (Gudmundsson and Sambridge, 1998) are drawn by magenta lines at 50 km intervals

15 from the trench.

17 Fig. 4: Interpolated CMP gathers along profiles A to D (see Fig. 1A) superimposed on high-

18 resolution P-wave velocities (Obayashi et al., 2006). Also indicated are earthquakes within the

19 averaging window of each cross-section. Our interpretations (white lines) are combined with the

20 -0.5% velocity perturbation contours (red lines, Obayashi et al., 2006).

22 Fig. 5: (A) Mantle reflectivity structures along the northernmost profile E (see Fig. 1B), P-wave

23 speeds (Obayashi et al., 2006), and Wadati-Benioff zone seismicity (yellow circles). The thin

52 1 red-lines outlines -0.5% velocity perturbations (Obayashi et al., 2006). Our interpretations are

2 highlighted by the dashed white lines. (B) A schematic interpretation of the HRZs for the Japan-

3 Kuril subduction system. The thick red line along the surface of the subducting slab indicates the

4 ongoing process of dehydration melting. Slab penetration is likely in regions where the 660

5 appears to be segmented.