Mid-Mantle Heterogeneities Associated with Izanagi Plate: Implications for Regional Mantle Viscosity ∗ Juan Li A, , David A

Earth and Planetary Science Letters 385 (2014) 137–144

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Earth and Planetary Science Letters

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Mid-mantle heterogeneities associated with Izanagi plate: Implications for regional mantle viscosity ∗ Juan Li a, , David A. Yuen b,c a Key Laboratory of the Earth’s Deep Interior, Institute of Geology and Geophysics, Chinese Academy of Sciences, 100029 Beijing, China b School of Environment Sciences, China University of Geosciences, 430074 Wuhan, China c Department of Earth Sciences, University of Minnesota, Minneapolis, MN 55455, USA article info abstract

Article history: We employed S-to-P converted waveforms to detect mid-mantle scattering beneath northeastern China Received 20 July 2013 and the adjacent Japan Sea. Broadband and short-period waveforms recorded by seismic arrays for eight Received in revised form 16 October 2013 moderate-sized deep earthquakes were analyzed using a non-linear array stacking technique, and mid- Accepted 23 October 2013 mantle scatterers within the depth range 930–1120 km were clearly revealed. The heterogeneities have Available online 12 November 2013 an overall lateral extent of ∼800 km, and mostly occur within a region with a high velocity anomaly. Editor: P. Shearer The accumulation of MORB-like slab materials at mid-mantle depths might cause a different chemical Keywords: composition than that of the surrounding peridotitic mantle. The spatial isolation of the heterogeneities mid-mantle discontinuity from the stagnant Paciﬁc slab suggests an origin related to the subduction of ancient Izanagi plate. subduction In combination with the reconstruction history of plate motions, we estimate the viscosity of the topmost Izanagi plate lower mantle to vary from 1.0 × 1022 to 1.6 × 1023 Pa s, which can be used as an independent constraint viscosity on the rheology of the lower mantle on a regional scale. © 2013 Elsevier B.V. All rights reserved.

1. Introduction depth of ∼920 km beneath the Tonga subduction zone (Kawakatsu and Niu, 1994), and scatterers around depths of 1000–1200 km Mapping the mantle heterogeneities in the lower mantle can and 1600 km have been mapped beneath Izu–Bonin–Mariana and assist our understanding of the distribution of geochemical reser- other subduction zones (Niu and Kawakatsu, 1997; Kaneshima and voirs and the scale of mantle circulation (e.g., Kawakatsu and Helffrich, 1999; Castle and Creager, 1999; Vinnik et al., 2001; Niu, 1994; Kaneshima and Helffrich, 1999; Helffrich and Wood, Vanacore et al., 2006; Niu, 2013). A velocity discontinuity at 2001; Castle and van der Hilst, 2003). The sources and loca- ∼1050 km has been revealed beneath mantle upwelling regions, tions of the distinct chemical reservoirs that provide the differ- e.g., Iceland and the Hawaiian Islands, arguing against a globally ent signatures of ocean island basalts (OIB) and mid-mantle ridge continuous mineralogical phase change near the depth of the mid- basalts (MORB) has been actively debated (e.g. Hofmann, 1997; mantle (Shen et al., 2003). Kellogg et al., 1999). Models with reservoir boundaries at different Seismic tomography has revealed a prominent stagnant slab depths in the mantle and chemically distinct blobs embedded in lying horizontally in the upper mantle transition zone (MTZ) un- the lower mantle have both been invoked to explain the geochem- der the Japanese subduction zone and extending 1000–2000 km ical, heat flow and seismic observations (Wen and Anderson, 1997; to the west (Fukao et al., 2001; Huang and Zhao, 2006), which Kellogg et al., 1999; Tackley, 2000). makes this location an ideal place to investigate geodynamical Compared with the heterogeneous upper mantle, the lower problems (Fig. 1). We undertook a systematic and thorough search mantle seems to be generally homogeneous, except for the lower- for deep mantle heterogeneities beneath northeastern China and most several hundred kilometers above the core–mantle boundary. the adjacent Japan Sea. Array stacking techniques were applied A smoothly increasing wave speed and density profile through to detect the weak signature of the scattered waves indicative the lower mantle has been inferred from various global mod- of the lower mantle structure. Broad mid-mantle scatterers with els (Dziewonski and Anderson, 1981; Kennett and Engdahl, 1991). depths ranging from 930 to 1120 km were revealed to the east Studies using scattered seismic waves, however, have detected seis- of the trapped stagnant slab. We compared our results to deep mic discontinuities and/or reflectors in the mid- to lower mantle. images obtained from seismic investigations and plate reconstruc- A local discontinuity with little topography has been imaged at a tion studies. We argue that subduction of the ancient Izanagi plate has caused the mid-mantle scatterers. From the spatial distribu- * Corresponding author. tion of this heterogeneity, the local rheology of the top portion E-mail address: [email protected] (J. Li). of the lower mantle can be constrained quite well, which in turn

Fig. 1. Location of the eight moderate-sized events investigated and the station distributions. (a) The large beachballs show the source mechanisms of the seven events with the S1000P phase observed; the pink dot indicates event #8 with no S1000P identified. The red point clusters are the conversion points calculated at their estimated depth for the S1000P phase, and the adjacent numbers indicate the depths of the observed scatterers. The P velocity model (Fukao et al., 2001) for the depth range 900 to 1000 km is shown in the background. Line AB indicates the position of the cross-section shown in Fig. 5. Yellow points indicate the conversion location of SdP phase of the Mw 6.9 earthquake analyzed by Niu (2013). (b) Distribution of seismic arrays; the triangles indicate the stations used. Insert is a schematic plot showing the ray paths of the P and SdP phases. provides an additional constraint for unraveling of the dynamical 2.1. Preprocessing processes and tectonic history under the western Pacific up to at least 50 Myr. We processed the raw data using the following basic steps. First, we examined every vertical-component seismogram and excluded bad traces and those with SNR < 5. When necessary, we re- 2. Data and method versed the polarity of the recorded seismograms. We applied a band-pass filter between 1 and 5 s to the retained seismograms, and handpicked the first peak of the P waves. We then aligned We used source-sided S-to-P conversion waves to detect mid- these to zero seconds, which was then used as a reference time mantle scatterers. This kind of wave, named SdP, starts as a down- for further stacking. We normalized the seismograms according to going S wave and is subsequently converted to a P wave at the their maximum amplitude in the time window −10 to 120 s rel- mantle discontinuity or reflector (Fig. 1). Compared to the direct ative to the direct P waves (Kawakatsu and Niu, 1994; Castle and P wave, the mid-mantle S-to-P converted wave has a lower slow- Creager, 1999). A shorter time window was selected for two earlier ness, and is thus received at a steeper angle of incidence. Because seismic records because of the limitations of the record time. of the small amplitude of the SdP phase, a large-scale regional The waveforms recorded by UW for events #2 and #7 are seismic array stacking technique (e.g., Kawakatsu and Niu, 1994; shown in Fig. 2(a)–(b). As well as the direct P and surface con- Li et al., 2008) was used to enhance the coherent later arrivals, verted pP phases, we found another clear phase ∼40–50 s after the which helped to identify the mid-mantle discontinuities. direct P wave in most of the individual seismograms. No seismic We studied a total of eight moderate-sized deep earthquakes waves would be expected to arrive in this time window according (5.3 mb 6.0) that occurred after the year 1980 beneath the to the 1D global reference model. The measured particle motion, Russian–Chinese border and the Japan Sea (Fig. 1(a)). For earth- the incident angle, the approaching direction of this arrival, and quakes that occurred before 2008, we used the EHB catalog with later detailed analysis suggest that this is an S to P conversion the source parameters relocated (Engdahl et al., 1998); for later wave at a mid-mantle depth of ∼1000 km (named as S1000P here- events, the location information was taken either from the NECI after). No S to P converted phase is visible in any individual records catalog or previous work, if available (e.g., Li et al., 2013). All these in a similar time window for the other six events. events have a simple source-time function, which significantly re- duces contamination of the complex source rupture process to the 2.2. Nth-root slant stacking weak later arrivals. The seismic arrays used in this study include the Pacific Northwest Regional Seismic Network (UW), the Caltech We applied Nth-root stacking to enhance the low amplitude of Regional Seismic Network (CI), the German Regional Seismic Net- S to P conversions from mantle discontinuities. Details of the stack- work (GRSN) and the Alaska Regional Network (AK) (Fig. 1(b)). ing procedure are given by Kawakatsu and Niu (1994) and Li et al. Both the UW and CI networks have ∼100–200 short-period sta- (2008). We varied the slowness within the range ±1s/deg (rela- tions, while the GRSN and AK arrays consist of ∼30 and ∼20 tive to P) with an increment of 0.02 s/deg. The time correction for broadband seismic stations, respectively. each trace was calculated using a plane wave assumption before J. Li, D.A. Yuen / Earth and Planetary Science Letters 385 (2014) 137–144 139

Fig. 2. Aligned seismograms recorded by the UW array for event #2 (a) and #7 (b) with clear phases identified in individual records in the time windows ∼40–50 s and ∼45–55 s after the direct P wave, respectively. No such phase is observable in the individual records for the other events. each stacking. Using the results, we produced a vespagram, which during the initial radiation process, we expected a comparable am- is a diagram showing seismic energy as function of slowness and plitude for both phases before the S-to-P conversion. We calculated time. The ‘hotter’ color clusters (Fig. 3) represent a greater energy the theoretical amplitude of S660P at the 660-km discontinuity for of the possible phases when the assumed slowness is consistent event #8 based on the iasp91 model for an assumed UW array. The with the incoming apparent velocity. In the stacked seismograms, value is ∼0.006, which is rather lower than that of ∼0.2–0.3 cal- the mid-mantle S to P converted phase is characterized by a neg- culated for other event–array pairs, with the S1000P phase clearly ative slowness and an arrival time between the P and pP phase identified. We consider that the lack of S1000P arrival for event #8 arrivals. More specifically, it will arrive a little later than the S660P, can be explained by an unfavorable focal mechanism. which is the converted wave at the 660-km discontinuity. The amplitude of the nth root slant stacking waveforms of phase S1000P is in general 4–6 percent of the direct P waves. For 2.3. Azimuth–slowness stacking event #2, however, the energy of the stacked deep mantle phase is a little higher, with values ∼7 and 12 percent that of the di- Slant stacking method is based on the assumption that the rect P wave for the UW and CI array stacking, respectively (Fig. 3). laterphasearrivesfromthesameraypathofthedirectPwave. The amplitude of S-to-P conversion depends linearly on the shear Non-horizontal and lateral heterogeneous structure will gener- velocity jump across the discontinuity and is insensitive to the ate phases that propagate out of the great circle path. We use compressional velocity and density (Castle and Creager, 1999). The the array beam-forming technique (e.g. Aki and Richards, 1980; observed peak amplitudes of the S P wave ranged from 40% Rost and Thomas, 2002) to measure the back-azimuth deviations 1000 to 120% of S660P. This means that if the conversion occurs at a of the S1000P phases relative to the great circle path. Waveforms sharp boundary, then an S-wave velocity increase of approximately within 3 s around the S1000P phase envelope peak are linearly 40–120% of that of the 660-km discontinuity, or 0.4–1.2km/s, stacked, and the energy of the stacked seismogram within the time window is calculated. The increments of azimuth and slowness are is required to generate the observed amplitude of the S1000P phase. ◦ taken as 0.2 and 0.1s/deg in the slowness–azimuth stacking, re- Fig. 4 shows azimuth versus slowness stacking results. To illus- spectively. trate the relative amplitude between S1000P phase and the background noise, a linear scale is used in calculating the stacked 3. Results energy. The intersection of the two straight lines is the expected location of the direct P wave. Generally the azimuthal and slow- ◦ In general, second root slant stacking (Li et al., 2008) was ap- ness resolution for UW array are about 0.8 and 0.1 s/deg if taken plied, and if necessary, a higher root stacking was adopted to en- the value of 0.8 as the uncertainty estimation level (Kaneshima hance the coherent weak later arrivals. In Fig. 3, the mid-mantle and Helffrich, 1998); while the azimuth resolution for CI appear conversion phase S1000P can be clearly identified in the stacking a little poorer (Fig. 4). The slowness measured for the later S- vespagrams for seven of the eight events shown. Because S660P to-P converted phases is slightly smaller than that of the direct and S1000P are characterized by slightly different takeoff angles P wave (Fig. 4), and the arrival azimuths of later waves deviate 140 J. Li, D.A. Yuen / Earth and Planetary Science Letters 385 (2014) 137–144

Fig. 3. Color contour map of the second-root stacked traces of all the event–array pairs with the S1000P phase clearly identiﬁed. “Hotter” color clusters represent greater energy and indicate a phase arrival. S1000P arrivals, which are identiﬁed by the slowness and arrival time, are indicated by black arrows. The conversion phase at the shallower upper mantle discontinuities, where detectable, is marked.

◦ fromthedirectPwavebylessthan1.0 (Table 1), suggesting a arrival after the direct P wave. Two deep earthquakes are used and sub-horizontal structure beneath those sources. one of them is same as event #2. The estimated arrival direction ◦ We used the iasp91 velocity model (Kennett and Engdahl, 1991) of this event to the array is 0.5 from the great circle ray path, ◦ to convert the differential travel time between S1000P and P to the which is consistent with our measurements (0.6 for CI and UW conversion point depth (Fig. 5). We measured the relative travel array). The estimated S to P conversion depths is 1043 km, which time difference accurately from the linear slant stacks, which pre- is same as our estimation from CI array, and a little shallower than served the shape of waveform, for all available event–array pairs. the depth estimated from UW and AK array (Table 1). This is rea- We also calculated the relative slowness for an S–P wave converted sonable because of the large array aperture used in his study, and at the observed depth for each event (Table 1). A horizontal dis- our ﬁner results revealed the slight difference of the location of continuity was assumed in the calculation. The observed slowness scatterers from different azimuths. is in general quite consistent with the predicted value, with the The depths of the observed mid-mantle scatterers beneath largest discrepancy being ∼0.07 s/deg (Table 1). northeastern China and Japan Sea are by no means uniform but Recently, Niu (2013) has applied array stacking techniques for vary in the range 930–1120 km (Fig. 5). The observations within ◦ the USArray data to reveal the nature of a clear abnormal seismic the longitudinal range 135–137 E are a little sparse, and there is J. Li, D.A. Yuen / Earth and Planetary Science Letters 385 (2014) 137–144 141

Fig. 4. Azimuth versus slowness stacking in a time window around the P phase (a) and S1000P phase (b–d). (b)–(c): Event #2 for UW and CI array; (d): event #7 for UW array. A linear scale is used in calculating the energy to illustrate the amplitude between the S1000P phase and the background noise level.

Table 1 Source parameters used in this study and observations of the S1000P.

Event Origin date Lat. Lon. Mag. H Array Az T obs pobs pcal Depth ◦ ◦ ◦ # ( ) ( ) (mb) (km) ( ) (s) (s/deg) (s/deg) (km) 1 05/10/2011 43.29 130.94 5.4 544.0b CI – 37.50 −0.18 −0.13 1002 2 05/19/2008 42.50 131.87 5.7 519.0c UW 0.6 45.20 −0.24 −0.18 1050 AK – 50.85 −0.28 −0.33 1055 CI 0.6 43.00 −0.22 −0.15 1043 3 03/09/2007 43.23 133.54 6.0 441.2 CI – 44.85 −0.10 −0.14 970 4 02/13/2000 42.80 131.67 5.4 515.7 GRSN – 35.7 −0.08 −0.12 937 5 08/11/1992 43.28 135.54 5.3 363.4 UW – 52.25 −0.18 −0.19 943 6 05/11/1990 41.79 130.95 5.7 578.0 UW – 35.50 −0.18 −0.14 1001 CI – 34.40 −0.12 −0.12 1004 7 02/11/1987 43.12 132.23 5.5 504.0 UW 1.0 51.95 −0.28 −0.22 1114 8a 10/08/1983 44.17 130.79 5.6 563.8 UW – – – – –

Abbreviations are: Az: azimuth relative to direct P wave from beamforming technique; pobs: slowness measured from the vespagram; pcal: theoretical slowness for the S1000P phase calculated from iasp91 model; T obs: observed arrival time measured from the stacked vespagram. a Not preferred source mechanism. b Depth taken from NEIC. c Depth taken from Li et al. (2013). only one point with depth ∼970 km in the middle of this area. ( Karato and Karki, 2001). We calculated the time correction for A mid-mantle discontinuity with depth ∼943 km is clearly iden- the shallowest and deepest events; the largest variations for P and tifiable at the east end of profile AB (Fig. 4). The lateral extent of S1000P waves were −1.85 s and −1.50 s, respectively. The relative the mid-mantle scatters is overall wide, with a value of ∼800 km, travel time between the two phases, however, varied by no more arguing for a consistent local structure and not a point source. than 0.4 s, which corresponds to a depth variation of 5 km. Differ- ent tomographic models, e.g., Fukao et al. (2001) were applied, and 4. Discussion and conclusions the variation in the relative time difference between the direct P We estimated the influence of 3D structure on the relative and S1000P phases is consistently trivial, with an absolute value of ∼0.2–0.4 s. In addition to the uncertainty in hypocentral location travel times between the direct P phase and the S1000P phase to account for the effects of the complex structure on the time–depth and picking errors, the uncertainty in depth measurement should conversion (Li et al., 2008). We used the shear wave velocity model be no more than 10 km. of Grand (2002) and a P wave velocity model converted from the Mid-mantle discontinuities/scatterers with depths between 900 shear model with a scaling relationship of d ln V s/d ln V p ∼ 1.6 and 1700 km have been identified in a wide range of different 142 J. Li, D.A. Yuen / Earth and Planetary Science Letters 385 (2014) 137–144

ers reflects a compositional anomaly associated with subducted basaltic crust (MORB) in the deep mantle (e.g. Kaneshima and Helffrich, 1998; Kaneshima and Helffrich, 1999; Castle and van der Hilst, 2003; Niu, 2013). There are significant differences between the MORB and the pyrolite mantle under lower mantle conditions (Hirose et al., 1999; Hirose et al., 2005; Irifune et al., 2010). The high content of alu- minum in MORB will expand the region of stability of garnet to a depth of around 800 km, which is ∼70 km deeper than that of the peridotitic mantle. Below this, the MORB will predominantly be composed of perovskite and Ca-perovskite, aluminous phase with the Ca-ferrite structure, and stishovite (Kingma et al., 1995; Hirose et al., 1999); whereas in a peridotitic mantle, it will be mostly composed of perovskite, Ca-perovskite and magnesiowüstite (Ono et al., 2001). The presence of substantial amounts of MgO in the peridotitic composition, and the aluminous phase and stishovite in MORB, thus give rise to substantial differences in elasticity, which can be observed using short-period seismic waves. After the transformation of MORB to the perovskite lithology at a depth of ∼720 km, the basaltic MORB becomes denser than the surround- Fig. 5. Depth cross section along line AB in Fig. 1 with the conversion points of ing pyrolitic mantle, and then may sink gravitationally into the S1000P indicated by dense red dots. The depth section of the P wave velocity model deep mantle (Hirose et al., 1999, 2005). (Fukao et al., 2001) is also plotted, with blue and red colors indicating fast and slow To constrain the source of the subducted basaltic crust stalling velocity anomalies. The black dots mark the undulation of the 660-km discontinuity investigated by S-to-P scattered wave (Li et al., 2008), and the rapid change of the in the mid-mantle depths, we compiled recent seismic studies of 660-km discontinuity constrains the hinge of the slab. Yellow dots are location of the deep structure beneath the Japan Sea and northeastern China. conversion points of SdP phase for the Mw 6.9 event analyzed by Niu (2013). A rapid change in depth of the 660-km discontinuity near longi- ◦ tude 131 E constrains the spatial location of the slab hinge at a tectonic environments, including subduction zones (e.g. Niu and fine scale (Fig. 5), indicating where the lower boundary of the sub- Kawakatsu, 1997; Castle and Creager, 1999), mantle upwelling re- ducting Pacific lithosphere encounters the 660-km discontinuity (Li gions (Shen et al., 2003) and the African rift region (LeStunff et al., 2008). We noticed that all of the observed mid-mantle scat- et al., 1995). The geometry of the overall mid-mantle scatter- terers, seem to locate east of the hinge of the stagnant slab (Fig. 4). ers over a 200 km depth range seen in our study appears com- To check possibility of a scatterer locating west of the slab hinge plicated; they are either tilted or strongly deformed, with large or beneath the stagnant part of the slab, we reexamined the other variationindepth.Itisdifficulttoexplainthedepthvariabil- deep earthquakes occurred in the study region. We find that de- ity in terms of a global mineralogical phase transition unless a spite 3 large earthquakes with Mw ∼7untiltheyear2011,only huge lateral variation in temperature is present. In addition, re- small events with Mw < 5.0 occurred several or tens kilometers cent mineral physics experiments have shown that, at the con- deeper than event #6 – the deepest one used in our study (Ta- ditions found in the top of the lower mantle at a pressure of ble 1). Their spatial distribution allows for a check for a possible ∼36 GPa, no known phase transition can be directly responsible scatterer west of the hinge, however, the energy released by them for the seismic discontinuity observed at a depth of ∼920 km in is very limited and inhibit this kind of scatterer detection. A large either the peridotitic mantle or in the MORB (Ono et al., 2001; event (Mw 6.9) with epicenter depth ∼577.7 km (NEIC location: ◦ ◦ Hirose et al., 2005). 42.586 N, 130.703 E) was analyzed in Niu’s (2013) recent work, The electronic spin transition of Fe occurs in (Mg, Fe)O fer- which is excluded in our study due to the complexity of its source ropericlase at pressures of 40–70 GPa in the lower mantle (e.g., time function. The estimated S-to-P conversion depth is 987 km Badro et al., 2003). The high-to-low spin transition might have (Niu, 2013). We calculated the conversion points according to the a significant effect on the single crystal elasticity of ferroperi- array geometry used in his study (Fig. 1) and projected them along clase, however, no obvious deviation in the aggregate seismic ve- the profile AB (Fig. 5). They are all located east of the slab hinge locities can be detected for the conditions in the depth range without any exclusion. It is thus reasonable to assume that the de- ∼1000–1800 km (Antonangeli et al., 2011). tected scatterers are located in the Pacific side of the hinge of the Seismic scattered wave has been applied to reveal undulation stagnant slab, which further suggests that these heterogeneities are of the mantle boundaries and the hinge of the subducting slab not directly connected with the ongoing subduction of the Pacific and the 660-km discontinuity beneath the study region has been plate. Alternatively, the ancient Izanagi plate subduction is one of well constrained (Li et al., 2008). A wide region of slow velocity the possible candidates for being the cause of the heterogeneities. is found to lie behind and below the hinge (Fukao et al., 2001; The exact date for the formation of the Izanagi plate remains Zhao et al., 2011). Despite the low resolution of the tomographic uncertain, although it is widely accepted that the Izanagi plate ex- map around this depth, the slow wave anomalies are observed isted for a long time before the birth of the Pacific plate (Woods to extend deep into the lower mantle. We calculated the loca- and Davies, 1982). By incorporating the continued spreading of tion of the conversion points for each event–array pair based on the Izanagi–Pacific ridge, Whittaker et al. (2007) proposed that the estimated conversion depth (Fig. 5). We noticed that except a mid-ocean ridge subduction beneath southern Japan occurred for one point at the west end, which is located at the boundary at 60–55 Ma. Seton et al. (2012) reconstructed the ocean basin between the fast and slow velocity regions, all the other con- evolution up to 200 Ma based on updated and more accurate mag- version points are located within a weak high velocity region, netic anomaly timescales, and argued that the disappearance of the separated by the pervasive low velocity anomalies (Fukao et al., Izanagi–Pacific ridge occurred around 55–50 Ma. 2001) without exception. Based on the close relationship between Taking the time of cessation of spreading between the Izanagi the observed largely deformed mid-mantle scatterers and the high and Pacific plates as a minimum estimate for the age of the ob- velocity anomaly, we argue that the observed mid-mantle scatter- served mid-mantle heterogeneities, we can estimate the settling J. Li, D.A. Yuen / Earth and Planetary Science Letters 385 (2014) 137–144 143

Fig. 6. Schematic diagram showing the location of the mid-mantle heterogeneity, which might be related to the ancient Izanagi plate subduction. Time of transportation of Izanagi plate related anomalies from the surface to the observed depth can be estimated from history of plate reconstruction and our seismic observations. Assuming those scatterers (thick red bar) roughly constraining the spatial distribution of the heterogeneities, the viscosity of the topmost lower mantle beneath the study region can be estimated.Thepurplelineisthetopographyofthe660-kmdiscontinuity(Li et al., 2008) revealed by seismic scattered waves.

22 23 time ts of the ancient sinking slab to first order, assuming a typ- region to lie within the range 1.0 × 10 to 1.6 × 10 Pa s, which ical scale of ∼200 × 800 km constrained by the spatial distribu- can be used as an independent constraint on the rheology of the tion of mid-mantle heterogeneities (Fig. 6). Based on two sim- lower mantle beneath the northwestern Pacific subduction zone. plified but reasonable assumptions: an average subduction rate ◦ of ∼10 cm/a and a dip angle of ∼45 (Engebretso et al., 1992), Acknowledgements it took 10 Myr to subduct ∼1000 km of slab to the base of the upper mantle (Fig. 6). Adding another 10–20 Myr of pile or We thank Satoshi Kaneshima and an anonymous referee for stagnation before sufficient negative thermal buoyancy was accu- valuable comments, which greatly improve the manuscript. Thanks mulated, means that the slab was 20–30 Myr old, when it be- are due to Gabriele Morra and Steven Jacobsen for some stimulat- gan to sink from the 660-km discontinuity. Then the slab would ing discussions. J. Li wants to thank Fenglin Niu for the inspired have sunk more or less vertically in the lower mantle. The av- thoughts and talks. The manuscript is finished while J. Li visiting erage settling velocity vs in the lower mantle is estimated to CU, Boulder, and she gratefully acknowledges the support of K.C. be 1.1–2.0 cm/a for sinking a distance of ∼400 km to the ob- Wong Education Foundation, Hong Kong. We also thank IRIS and served mid-mantle depth during this second period of 20–35 Myr. GRSN data centers for providing waveform data. GMT software was It should be noted that this result was obtained from independent used for plotting most of the figures. This research was supported geophysical inferences based on the history of plate reconstruc- by NSFC (J. Li, Grants 41074034 and 41274065) and NSF geochem- tion and seismic information, and thus can be used as a constraint istry programs. on the rheology of the lower mantle regionally. We used the vis- 2 cosity renormalization factor 2/9 × ρ gR /vs (Morra et al., 2010; References Capitanio et al., 2007)fromStokeslaw,whereρ is the differential density, taken to be in the range 1–2% (∼50 kg/m3); g is the Aki, K., Richards, P.G., 1980. Quantitative Seismology: Theory and Methods. W.H. Freeman, New York. 932 pp. gravity constant; and R is the effective length of the slab with R = ×[ + ] Antonangeli, D., Siebert, J., Aracne, C.M., Farber, D.L., Bosak, A., Hoesch, M., Krisch, b 1 ln(a/b) , where a and b are the longest and second longest M., Ryerson, F.J., Fiquet, G., Badro, J., 2011. Spin crossover in ferropericlase semi-axis of the sinking slab, here restricted to be 400 and 100 km, at high pressure: a seismologically transparent transition? Science 331, 64–67, respectively (Fig. 6). 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