Anomalously low amplitude of S waves produced by the 3D structures in the lower mantle

Physics of the Earth and Planetary Interiors 256 (2016) 26–36

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Anomalously low amplitude of S waves produced by the 3D structures in the lower mantle Akiko To a,⇑, Yann Capdeville b, Barbara Romanowicz c,d,e a

Japan Agency for Marine-Earth Science and Technology, 3173-25 Showa-machi, Kanazawa-ku Yokohama, Kanagawa 236-0001, Japan Laboratoire de Planétologie et Géodynamique de Nantes, CNRS, Université de Nantes, Nantes, France c Berkeley Seismological Laboratory, UC Berkeley, Berkeley, United States d Collège de France, Paris, France e Institut de Physique du Globe de Paris, Paris, France b

a r t i c l e

i n f o

Article history: Received 25 December 2015 Received in revised form 14 April 2016 Accepted 14 April 2016 Available online 23 April 2016 Keywords: Seismic wave propagation ULVZ LLSVP Core-mantle boundary Amplitude anomaly Spectral element method

a b s t r a c t Direct S and Sdiff phases with anomalously low amplitudes are recorded for the earthquakes in Papua New Guinea by seismographs in northern America. According to the prediction by a standard 1D model, the amplitudes are the lowest at stations in southern California, at a distance and azimuth of around 95° and 55°, respectively, from the earthquake. The amplitude anomaly is more prominent at frequencies higher than 0.03 Hz. We checked and ruled out the possibility of the anomalies appearing because of the errors in the focal mechanism used in the reference synthetic waveform calculations. The observed anomaly distribution changes drastically with a relatively small shift in the location of the earthquake. The observations indicate that the amplitude reduction is likely due to the 3D shear velocity (Vs) structure, which deflects the wave energy away from the original ray paths. Moreover, some previous studies suggested that some of the S and Sdiff phases in our dataset are followed by a prominent postcursor and show a large travel time delay, which was explained by placing a large ultra-low velocity zone (ULVZ) located on the core-mantle boundary southwest of Hawaii. In this study, we evaluated the extent of amplitude anomalies that can be explained by the lower mantle structures in the existing models, including the previously proposed ULVZ. In addition, we modified and tested some models and searched for the possible causes of low amplitudes. Full 3D synthetic waveforms were calculated and compared with the observations. Our results show that while the existing models explain the trends of the observed amplitude anomalies, the size of such anomalies remain under-predicted especially at large distances. Adding a low velocity zone, which is spatially larger and has less Vs reduction than ULVZ, on the southwest side of ULVZ, contributes to explain the low amplitudes observed at distances larger than 100° from the earthquake. The newly proposed low velocity zone mostly overlaps with the northern part of the Pacific large low shear velocity province (LLSVP) revealed in tomographic models. Although the very low amplitudes observed at a distance of about 95° remain unexplained, our results indicate that the boundary of the Pacific LLSVP is sharp, and the amplitude of S waves at these large distances is lowered by strong vertical and/or lateral deflection at the boundary toward the interior of the low velocity province. Ó 2016 Elsevier B.V. All rights reserved.

1. Introduction Global shear velocity (Vs) tomographic models have revealed large-scale low velocity structures at the base of the mantle beneath the Pacific and Africa, which commonly appear as strong degree-2 pattern in global tomographic models (e.g., Takeuchi, 2007; Houser et al., 2008; Kustowski et al., 2008; Ritsema et al.,

⇑ Corresponding author. Tel.: +81 45 778 5596; fax: +81 45 778 5439. E-mail address: [email protected] (A. To). http://dx.doi.org/10.1016/j.pepi.2016.04.001 0031-9201/Ó 2016 Elsevier B.V. All rights reserved.

2011; French and Romanowicz, 2014). On the other hand, forward modelling of shorter period seismic waveforms and travel times have documented the presence of ultra low velocity zones (ULVZs), which have larger velocity reductions and are spatially more localised (<1000 km; Cottaar and Romanowicz, 2012; Thorne et al., 2013a) relative to the degree-2 feature, near the core-mantle boundary (CMB; e.g., Garnero et al., 1998; Lay and Garnero, 2011). Resolving the spatial relation between the ULVZs and large-scale low Vs structures may hold the key to understanding the upwelling from global mantle circulation (e.g., Garnero and McNamara, 2008).

A. To et al. / Physics of the Earth and Planetary Interiors 256 (2016) 26–36

Many forward modelling studies of detailed seismic structures in the lowermost mantle are based on the analysis of travel time anomalies and waveform distortions (e.g., Ni et al., 2002; To et al., 2005; He and Wen, 2009; Idehara, 2011; To et al., 2011; Cottaar and Romanowicz, 2012; Thorne et al., 2013a; Pachhai et al., 2015; Zhao et al., 2015). A small number of studies have quantitatively examined the absolute amplitudes of CMBsensitive phases and their sensitivity to these structures (e.g., Zhang et al., 2009; Rondenay et al., 2010; Thorne et al., 2013b). The western part of the lowermost mantle beneath Hawaii has been one of the target regions for the forward waveform modelling. Waveforms of the S and Sdiff phases (hereafter S/Sdiff) recorded in the southwestern U.S. for earthquakes in the Papua New Guinea region are heavily distorted by low Vs anomalies beneath the Pacific (To and Romanowicz, 2009; To et al., 2011 [hereafter TFT2011]; Cottaar and Romanowicz, 2012 [hereafter CR2012]). Previous studies have mostly focused on the distortion of waveforms and the identification of plausible structural models to explain the emergence of prominent postcursors to S/Sdiff waves. TFT2011 looked at the 2D effect of slow anomalies and showed that they could be explained by a ULVZ (dVs/Vs =
25%; width = 500 km; thickness = 80 km) located southwest of Hawaii. Later, US array stations were deployed and enabled a dense sampling of the lower mantle structure. CR2012 conducted 3D modelling and concluded that a thinner, horizontally large ULVZ (dVs/Vs =
20%; width 910 = km; thickness = 20 km) located southwest of Hawaii, could account for the emergence of postcursors. In addition to the ULVZ, TFT2011 proposed that, based on observations of two postcursors at high frequency (0.15 Hz), a broader low velocity zone (LVZ) with smaller Vs reduction compared to the ULVZ may exist, and partially or fully cover the ULVZ. Moreover, a recent tomographic model revealed concentrated low Vs anomalies embedded within the well known large degree-2 low Vs features at the base of the mantle. Some of the low Vs anomalies extend upwards as conduits, suggesting the existence of broad domelike plumes (French and Romanowicz, 2015). In this study, we focused on amplitude anomalies of S/Sdiff phases observed in the datasets of TFT2011 and CR2012. We first

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confirmed that these amplitude anomalies were mainly generated by lowermost mantle structures, and not by upper mantle structures or errors in focal mechanisms. The amplitude anomalies generated by the recently developed tomographic model SEMUCB_WM1 (French and Romanowicz, 2014) and the previously proposed ULVZ were evaluated. Finally, we examined whether the observed amplitude information included any signature of the LVZ or of the broad domelike plume structures adjacent to the ULVZ, and whether they contribute to resolve the Vs distribution in the lower mantle. 2. Observations The earthquake used in this study (hereafter Event 1) is a deep earthquake (depth = 414 km; Mw = 6.6) which occurred on 20 March 2010 in the New Ireland region of Papua New Guinea. Records were obtained by 650 stations in North America, including the stations of US-array (www.iris.edu/earthscope/usarray) through the Incorporated Research Institutions for Seismology Data Management Center (IRIS DMC). The configuration of earthquake location and stations is shown in Fig. 1. The S/Sdiff waveform data from this earthquke were also analysed by CR2012. In Fig. 2a and b, we show comparisons of observed and synthetic tangential component velocity waveforms of direct S waves at two selected stations, G05D (distance = 89.7°; azimuth = 44.8° from the earthquake) of the US-array, and DGR (distance = 92.3°; azimuth = 56.6° from the earthquake) of the Southern California Seismic Network (SCSN). Station locations are shown by coloured triangles in Fig. 1. Seismic instrument responses were deconvolved from the observed waveforms. Synthetic waveforms were calculated for the 1D Preliminary Reference Earth Model (PREM; Dziewonski and Anderson, 1981) using the Coupled-mode Spectral Element Method (Capdeville et al., 2003). The source mechanism was obtained from the Global centroid-moment-tensor catalogue (Global CMT; Ekström et al., 2012). The waveforms are shown in five different frequency ranges. Each trace was normalised by the maximum amplitude of the direct S phase of the PREM synthetic waveform, and no extra scaling is done. A comparison between Fig. 2a and b showed that the high frequency components of the

Fig. 1. Locations of earthquakes and seismic stations used in this study. The background model is an SH velocity model of SEMUCB-WM1 at the core-mantle boundary (CMB). The location of the ultra low velocity zone (ULVZ; thickness = 20 km; dVs/Vs =
20%) proposed by Cottaar and Romanowicz (2012) is shown by the solid line. The LVZ of Model 4 (thickness 100 km; dVs/Vs =
5%) is indicated by the solid grey line (for the relatively well constrained part) and the dashed grey line (part not constrained by the dataset). In Model 4, the region where ULVZ and LVZ overlap is filled by dVs/Vs of ULVZ (i.e.,
20%).

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S wave recorded at station DGR had much lower amplitudes than the synthetic waves. The observed amplitude is close to that of the PREM synthetic waveform at lower frequencies. An arrival time delay of 6 to 7 s was observed in the lower frequency component of the data. In contrast, amplitude anomalies were not observed at G05D, and only a travel time delay of less than a few seconds was apparent. The sS phase at station DGR also exhibits anomalously lowered amplitude, however part of it may be due to an error in the focal mechanism used in the PREM synthetic calculation, as further discuss in Appendix B. Low amplitude anomalies, such as the one observed at station DGR, were recorded across broad regions. Fig. 3 shows the spatial distribution of the observed amplitude and travel time anomalies of the S/Sdiff phase, which were defined as:

dAobs ¼ log10

Aobs APREM

dT obs ¼ T obs
T prem

!

ð1Þ ð2Þ

where Aobs and Aprem are the maximum amplitudes of the S/Sdiff phase measured from observed and PREM synthetic waveforms, respectively. The travel time anomaly dT obs was obtained by cross-correlating the observed and PREM synthetic waveforms. Measurements with a cross correlation coefficient of less than 0.8 were discarded. Anomalies were measured using broadband seismic waveforms Butterworth band-pass filtered in two different frequency ranges, one with corner frequencies at 0.01 and 0.03 Hz (hereafter 0.01–0.03) and the other with corner frequencies at 0.027 and 0.11 Hz (hereafter 0.027–0.11). In Fig. 3, the measurements were spatially averaged over an area corresponding to 1°  1° intervals in distance and azimuth from the earthquake to station, and plotted at the centre of each area. The spatially averaged measurements are hereafter denoted as dT obs and dAobs . The distribution of the original measured values, plotted at the station locations, is shown in Fig. S1. The number of the data points per grid used for spatial averaging is shown in Fig. S1c. A time window from 50 s before to 50 s after the synthetic arrival of S/Sdiff was used for both travel time and amplitude measurements. This time window was long enough to include the previously reported prominent postcursors to S/Sdiff, observed at stations in the southwestern U.S. (Fig. S2). The emergence of the postcursor lowers the cross-correlation coefficient between the observed and PREM synthetic waveforms and such measurements were discarded; therefore, the number of travel time measurements was small in the higher frequency range (Fig. 3b). The amplitude measured as the amplitude anomaly is the largest one between the main phase and the postcursors. In Fig. 4a, we marked the traces for which the amplitude of the postcursors was larger than that of the main phase by orange lines. In Fig. S2a and b, we

present the separately measured maximum amplitudes of the main phase and postcursors in the same way as in Fig. 3a. The amplitude ratio dAobs measured in the higher frequency range (0.027–0.11 Hz) was very low around southern California (Figs. 3a and 4a). The lowest amplitude ratio was measured at a distance of 95 and an azimuth of 55° from the earthquake and it gradually increases at stations away from this location. The arrival times of S/Sdiff were delayed toward the south (Figs. 3d and S3a), as previously reported by TFT2011 and CR2012. 3. Effect of focal mechanism errors on amplitude anomalies We examined whether errors in the focal mechanism used to calculate the PREM synthetic waveforms could cause the observed amplitude anomalies. We show that the strong frequency dependence of the observed amplitude anomalies dAobs could not be produced by perturbing the focal mechanism. Synthetic waveforms were computed for focal mechanisms whose strike, dip, and slip were perturbed by up to 10° with respect to the major double couple component of the original focal mechanism using the Direct Solution Method (Takeuchi et al., 1996). We measured the amplitude ratios between the synthetic waveforms computed using the perturbed and original focal mechanisms. In Fig. 5, the result for one of the perturbed mechanisms is plotted in the same way as in Fig. 4. The amplitude ratios measured in the two frequency ranges (0.01–0.03 Hz and 0.027–0.11 Hz) were almost identical, indicating that the small (i.e., 10°) changes in the focal mechanism produced no frequency dependence. We examined whether the Global CMT solution of Event 1 appropriately explained the observed amplitudes of the other major phases (e.g., depth phases or SS). The method is described in Appendix A. The result showed that the strike and dip of the preferred mechanism were shifted by 5° from those in the major double couple of the original mechanism (Fig. 5b). The preferred mechanism better explained the weak signal of depth phases (i.e., sS and sSS) at some stations, but did not significantly change the amplitude of the main S/Sdiff phases. Moreover, the distribution of the observed amplitude anomalies obtained using the preferred mechanism (Fig. S4a) was similar to the distribution obtained using the original focal mechanism (Fig. 3a). Based on these results, we used the original focal mechanism for all synthetic waveform calculations in this study. 4. Causal structures in the lower mantle Observations obtained for a different earthquake located close to Event 1 suggested that the low amplitude anomaly was likely

Fig. 2. Tangential component velocity waveforms filtered in five different frequency ranges (solid lines). The reference time is the expected direct S phase arrival for PREM. Each trace is normalised by the maximum amplitude of the PREM synthetic waveforms (grey thick line). (a) Observed waveform at station G05D of the US-array. (b) Observed waveform at station DGR of the Southern California Seismic Network. (c) Synthetic waveforms created by CSEM for Model 2. (d) Synthetic waveforms created by CSEM for Model 4. The travel time correction is added to the CSEM waveforms.

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Fig. 3. Distribution of spatially averaged observed amplitude and travel time anomalies. (a) Amplitude anomalies at 0.027–0.11 Hz. (b) Travel time anomalies at 0.027– 0.11 Hz. (c) Amplitude anomalies at 0.01–0.03 Hz. (d) Travel time anomalies at 0.01–0.03 Hz. The map is plotted in an azimuthal equidistance projection with the earthquake location placed at the centre. Larger views of this projection are shown in Figs. 6, 9, and 10.

caused by structures in the lower mantle. In Fig. 6, we show the distribution of the amplitude anomaly dAobs for an earthquake in Papua New Guinea, which occurred on 12 June 2003 (Mw h = 6.2; depth = 185 km; hereafter Event 2). Event 2 occurred before the deployment of the US-Array; therefore, the number of available stations was much lower than for Event 1. A comparison of Figs. 3a and 6 showed that the amplitude anomaly distributions were very different between Event 1 and Event 2, even though Event 2 occurred close to Event 1 (the horizontal distance between the two earthquakes was 400 km). The difference in amplitude distribution between the two earthquakes indicates that the anomalies were not caused by structures beneath the stations. Furthermore, Fig. 6 showed that low amplitude anomalies were also observed for Event 2; however, their location was shifted compared with those observed for Event 1. For Event 2, the lowest dAobs was observed around the state of Wyoming, while for Event 1, the lowest dAobs was observed around Southern California. In Fig. 6, we plotted the theoretical ray path of the 1D model from each earthquake towards the region of the lowest amplitude ratio. We found that the two arrows crossed at a distance of 30° from the earthquakes. For S waves that reached stations 100° from the earthquake, the ray-paths at a distance 30° were at depths of 2700 km (e.g., 200 km from the CMB). When estimated based on the Fresnel zone of such S waves, which becomes more than several degrees wide in the lowermost mantle, the causative structure

would be located within several hundred kilometres of the crossing point of the two ray paths; therefore, the cause of the amplitude anomaly was most likely located in the lower mantle. We also looked at waveforms of other deep and large earthquakes (Mw > 6.2; Depth > 150 km) in the Papua New Guinea region, including those on 21 July 2003, 17 October 2003, 5 August 2010, 7 February 2001, and 7 July 2013. However, the SH nodal planes of these earthquakes were directed to the azimuth of interest; therefore, stable amplitude measurements could not be made. Older earthquakes are recorded by fewer stations; however, an event on 2 May 1996 (Mw = 6.6; Depth = 512 km) was found to show anomalously low Sdiff amplitudes at station US.CBKS, located around the state of Kansas, U.S.A. The traces of this event were presented in Fig. 3 of TFT2011. The horizontal location of this earthquake was placed between those of Event 1 and Event 2. Furthermore, the direction of the ray-path that connected the earthquake and the station with the lowest Sdiff amplitude was consistent with those of Event 1 and Event 2; therefore, supports our conclusion that the causative structure exist in the lower mantle. 5. Modelling method We examined whether the observed amplitude anomalies could be explained by the wave deflection effect of the 3D elastic

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Fig. 4. Amplitude anomalies (dA) measured in two different frequency ranges. The lower frequency measurements (0.01–0.03 Hz) are shown in red, and the higher frequency measurements (0.027–0.11 Hz) are shown in blue. Each panel shows the measurements in 3° interval distances, as shown by the legend in the right column. (a) Spatially averaged measurements from observed waveforms. Orange lines mark azimuthal ranges where the amplitudes of the postcursors are larger than that of the main phase. (b) Model 1: Tomographic model SEMUCB_WM1. (c) Model 2: PREM + ULVZ of CR2012. (d) Model 3: Tomographic model + ULVZ of CR2012. (e) Tomographic model + ULVZ of CR2012 + enhanced slow region in the western Pacific (shown by the grey and grey dashed lines in Fig. 1).

lowermost mantle structures of existing Vs models and some modified models by conducting forward waveform modelling. Full 3D synthetic waveforms were calculated for each model and then compared with observed waveforms. The models were evaluated by computing the variance reductions in travel time and amplitude between the synthetic and observed waveforms of the S/Sdiff phase. The spatially averaged observed values (i.e., dT obs and dAobs ) in the distance range 83–125° and azimuthal range 35–64° from the earthquake were used for the comparison. The full 3D synthetic waveforms were computed using the Coupled-mode Spectral Element Method (CSEM, Capdeville et al., 2003), which couples a spectral element computation in the heterogeneous target region to a 1D normal mode computation for the rest of the Earth. In this study, the spectral element method was used to calculate wave propagation in the bottom 370 km of the mantle. Our dataset sampled the ULVZ proposed by CR2012, the height of which was constrained to be 20 km. To properly account for the effect of the 20 km thick ULVZ on the CMB, the height of the mesh element in the deepest mantle was set to 20 km. The 3D synthetic waveforms were calculated for up to 0.125 Hz. The variance reduction in travel time (VRtt) was defined as:

 rtt VRtt ¼ 1


rtt



3D 1D

 100

M  2 1X wi dT obs i M i¼1 M  2 1X ¼ wi dT obs
dT 3D i i M i¼1

ð3Þ

rtt 1D ¼

ð4Þ

rtt 3D

ð5Þ

dT 3D ¼ T 3D
T prem þ dT corr

ð6Þ

3D

where dT denotes 3D-synthetic travel time anomalies measured by cross-correlating 3D-synthetic and PREM synthetic waveforms, and dT corr describes the travel time corrections applied to the 3Dsynthetic travel times in order to compare them with observations. The corrections consist of ellipticity corrections (Kennett and Gudmundsson, 1996), altitude corrections, and corrections to account for the 3D heterogeneity in the mantle outside of the 370 km spherical shell, where wave propagation was calculated by a 1D normal mode summation in CSEM. The altitude corrections were obtained by dividing the station altitudes by 3.2 km/s. The corrections for mantle heterogeneity were calculated using ray theory from the SH velocity of tomographic model SEMUCB-WM1 (French and Romanowicz, 2014). The travel time corrections dT corr ranged between
0.9 and 7.3 s, which was not small compared with the observed travel time anomalies dT obs ranged between
2.8 and 10.4 s (Fig. S4a). The symbol wi denotes a weight applied to compensate for the differences in area used for spatial averaging of the original measurements. Since the area was defined by 1°  1° distance and azimuth from the earthquake, wi was dependent on the distance, and ranged from 1 at a distance of 90° to 0.82 at the furthest distance of 125°. M denotes the number of spatially averaged measurements. Although we measured travel time anomalies in five different frequency ranges (Fig. 2), here we only consider values measured in the frequency range 0.02–0.05 Hz, since this contained the largest number of measurements. The number of measurements decreased at higher frequency owing to the dissimilarity between observed and PREM synthetic waveforms,

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Fig. 5. (a) Ratios between synthetic S/Sdiff maximum amplitudes of a perturbed focal mechanism and the original Global CMT focal mechanism, are plotted in two frequency ranges. The structural model is PREM. The lower frequency measurements (0.01–0.03 Hz) are shown in red, whereas the higher frequency measurements (0.027–0.11 Hz) are shown in blue. The perturbed mechanism is the preferred mechanism shown in (b). (b) Left: The original focal mechanism obtained from the Global CMT catalogue. Right: The preferred mechanism, which better explains the amplitude ratios of observed data. (c) Left: Horizontal radiation patterns of SH at a take-off angle of 23.5° from the source (i.e., Sdiff towards a distance of 102° from the source). Solid lines indicate the range of azimuths represented in the dataset (39–65°). Right: SH radiation pattern in the vertical plane at an azimuth of 40.8°. (d) Comparisons between the observed and two synthetic waveforms calculated using the original and preferred focal mechanisms. The waveforms are filtered in the frequency range 0.02–0.05 Hz.

which was caused by the postcursors. Furthermore, it also decreased at lower frequency owing to the low signal to noise ratio.The variance reduction in amplitude (VRamp) was defined as:





ramp 3D  100 ramp 1D

ð7Þ

VRamp ¼

1


ramp 1D ¼

N  2 1X wi dAobs i N i¼1

ð8Þ

ramp 3D ¼

N  2 1X wi dAobs
dA3D i i N i¼1

ð9Þ

dA3D ¼ log10

A3D APREM

! ð10Þ

where dA3D denotes 3D-synthetic amplitude anomalies (i.e., the maximum amplitude of the S/Sdiff phase measured from 3Dsynthetics waveforms divided by the maximum amplitude of S/Sdiff phase obtained from the PREM synthetic waveforms). We used the amplitude measurements from five different frequency ranges indicated in Fig. 2 to evaluate the 3D models. N denotes the number of measurements, which was obtained by a summation of available measurements in every frequency range.

6. Results We examined approximately 30 models, including a tomographic model and the ULVZ of CR2012. Many of the modified models were created by placing a cylindrical LVZ on the CMB, with or without the tomographic model and UVLZ in the background. The shape was set to cylindrical in order to minimise the number of parameters that defined the LVZ. The radius, height, and Vs reductions of the LVZ were varied between 100–350 km, 450– 1000 km, and 3–5%, respectively. The horizontal location of the centre was changed between the distance and azimuth of 25–50° and 50–76° from the earthquake, respectively. The results obtained for four models are described in the following Sections 6.1–6.4, and are shown in Fig. 4 (amplitudes), Fig. S3 (travel times), and Table 1.

6.1. Tomographic model (Model 1) In Figs. 4b and S3b, we show the travel time and amplitude anomalies obtained for the tomographic model SEMUCB-WM1. While some observed features (Fig. 4a) at large distances were captured, including the increase in travel time delays toward the south (i.e., toward the large azimuths), amplitude anomalies were largely under estimated.

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Fig. 6. (a) Observed amplitude anomalies for Event 2. The amplitudes are measured from waveforms, band-pass filtered in the range 0.027–0.11 Hz. The locations of Event 1 and Event 2 are shown by the grey and light blue stars, respectively. The distances and azimuths from each earthquake are shown by the grey and light blue dashed lines with an interval of 10°. The grey and light blue arrows roughly indicate the direction from each earthquake to each region where the minimum amplitude ratio is observed. For Event 1, the region with the minimum amplitude ratio is located around a distance and azimuth of 95° and 55° from the earthquake, respectively. For Event 2, it is located around the distance and azimuth of 97° and 49° from the earthquake, respectively. (b) Synthetic amplitude anomalies for Event 2 calculated for Model 4.

6.2. PREM + the ULVZ of CR2012 (Model 2) In Figs. 4c and S3c, we show the amplitude and travel time anomalies obtained for the ULVZ structure proposed by CR2012. The location of the ULVZ is indicated by the black line in Fig. 1. The background model was PREM. Although the model proposed by CR2012 used a tomographic model in the background of the ULVZ (i.e., similar to our Model 3), here we solely examined the effect of the ULVZ on the observed amplitude. Fig. 4c shows that the frequency dependence of the observed amplitudes was captured to some extent, as documented by CR2012. At distances as short as 92°, the amplitudes in the higher frequency range were significantly decreased. The ray theoretical bottoming point of such a direct S phase for PREM was located 200 km above the CMB (Fig. S5); therefore, it should not be sensitive to the thin ULVZ in the high frequency range. In fact, the PREM synthetics S waveform around this distance is a superposition of the direct S and ScS waves, which arrived within 2 s of each other and had the same polarity. The ULVZ delay the ScS arrival separated it from the direct S wave; therefore, the amplitude of the S wave appeared smaller in Model 2 than in PREM. The 3D synthetic waveform (Fig. 2c, top trace) showed that the arrival time of the direct S wave was the same as that for PREM, whereas its amplitude was reduced. Furthermore, it showed a clearly delayed ScS after the direct S. At distances of greater than 100°, amplitudes in the higher frequency range contained a kink around the 57° azimuth (Fig. 4c). Table 1 Travel time (VRtt) and amplitude (VRamp) variance reductions of the four models described in Section 6.1–6.4.

VRtt VRamp

The 3D synthetic waveforms (e.g., Fig. 7) showed that beyond this azimuth, the amplitude of the postcursors became larger than the main phase. The amplitude of the main phase observed at azimuths of less than 57° were lower owing to wave deflection towards the south at the northern boundary of the ULVZ. This phenomenon can also be understood from the ray tracing conducted by CR2012 (Fig. 5 of CR2012), which showed how rays were laterally refracted by the ULVZ. The lower amplitude demonstrates that the wave amplitudes were indeed sensitive to the transverse gradients of an elastic structure (e.g., Woodhouse and Wong, 1986; Romanowicz, 1987; Romanowicz and Mitchell, 2015). 6.3. Tomographic model + the ULVZ (Model 3) In Figs. 4d and S3d, we show the amplitude and travel time anomalies obtained for the ULVZ structure of CR2012 using the tomographic model SEMUCB-WM1 in the background. The amplitudes measured for the high frequency waveforms were lower at distances around 101–103° (indicated by the thick orange line in Fig. 4d) compared with Model 2; therefore, this model explained the observation slightly better. The waveforms of this distance range (Fig. 7a) showed that the amplitudes of both the main phase and postcursors were reduced as compared with Model 2, with the amplitude reduction more significant for the postcursors than for the main phase. Our dataset sampled the northern large low shear velocity province (LLSVP), which is a feature in all global Vs models (e.g., Fig. 1; Lekic et al., 2012). The observed amplitude reduction in Model 3 was likely due to wave deflection toward the interior of the Pacific LLSVP, which is further discussed in Section 6.4.

Model 1 Tomographic model

Model 2 ULVZ of CR2012

Model 3 Tomographic model + ULVZ

Model 4 Tomographic model + ULVZ + LVZ

6.4. Tomographic model + the ULVZ + a LVZ between the event and the ULVZ (Model 4)

65.7 (%) 13.1

75.6 17.3

78.4 25.7

81.3 31.5

Our results showed that in addition to the ULVZ of CR2012 with the tomographic model in the background, placing a low velocity

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solid orange line in Fig. 9 and a solid grey line in Fig. 1. Its eastern end was set at an azimuth of 64° from the earthquake, which corresponds to the largest azimuth among the stations used in the modelling. The western end was set at a distance of 20° from the earthquake because the S/Sdiff phase of the distance range of our dataset (>83°) only has sensitivity beyond a distance of 20° for structures at the base of the mantle (see the figure caption of Fig. S6). The rest of the lateral boundary of the LVZ (plotted by a dashed line in Figs. 1 and 9) was not well constrained and was an artefact of setting the shape of the LVZ to cylindrical. The relatively well-constrained part of the LVZ (solid grey line in Fig. 1) was close to the northern border of the LLSVP of the tomographic model; therefore, Model 4 was almost equivalent to the structural model, which had a large Vs gradient at the boundary in the northern Pacific LLSVP. 6.5. Other models examined

Fig. 7. Observed waveforms (thick black lines) and 3D synthetic waveforms (black lines), obtained for models 1–4. All data were band-pass filtered at 0.027–0.11 Hz. Thick grey lines denote the PREM synthetic waveforms. The amplitude of each trace is normalised by the maximum amplitude of the PREM synthetic waveform. The travel time correction dTcorr is added to each 3D synesthetic waveform. (a) The observed waveform at station SC63 of the Seismic Investigation of Edge Driven Convection Associated with the Rio Grande Rift network (Pulliam et al., 2008), located at a distance and azimuth of 102.6° and 58.3° from the earthquake, respectively. (b) The observed waveform at station 529A of the US-array, located at a distance and azimuth of 102.2° and 59.9° from the earthquake, respectively.

zone (LVZ) at the base of the mantle on the southwestern side of the ULVZ contributed to further reductions in the observed variances. The LVZ is spatially larger and has a smaller Vs reduction than the ULVZ. In Figs. 4e and S3e, we show the amplitude and travel time anomalies for Model 4, which included an LVZ with a height and velocity reductions of 100 km and 5%, respectively. The location of the LVZ is shown by the grey circle in Fig. 1. The low amplitudes observed at the southern stations (i.e., those at a distance of 101–106°, as indicated by thick green line in Fig. 4e) were better explained by Model 4 than by Model 3. The waveforms in this distance range (Fig. 7b) showed that the amplitudes of both the main phase and postcursors were lower in Model 4 than in Model 3. The changes in amplitudes generated by the LVZ alone (Fig. 9) showed that amplitude reduction at the southern stations (Fig. 4e thick green line) was likely due to deflection of waves toward the interior of LVZ and a shadow zone where no theoretical ray arrives was produced. At this point, it remains unclear whether the lateral or vertical deflections played the dominant role, and this remains a subject for a future study. The portion of the lateral LVZ boundary that was constrained by our dataset was plotted using a

We show here the results for some of the other models examined, including those that explained the observed anomalies to the same degree as Model 4. We examined a series of models obtained by taking a contour line of the original tomographic model and filling the slower region with a constant Vs anomaly with respect to PREM (e.g., To et al., 2005). The saturated model was used from the CMB to 300 km above the CMB, then 3D Vs anomalies were gradually reduced to 0% from 300 km to 370 km above the CMB. For example, a model obtained by taking the
1% contour line of SEMUCB-WM1 and saturating the slower region with
3% Vs reduction with respect to PREM gave VRtt and VRamp of 71.0% and 27.8%, respectively. The saturated model had a larger Vs gradient than the original tomographic model at the LLSVP boundary. Since the saturated model contained similar structures to Model 4, it provided a similar level of variance reduction. In Fig. 10a, we show the results of a series of structures obtained by changing the horizontal location of a cylindrical LVZ placed on the CMB with the ULVZ of CR2012 in the background. The radius, height, and Vs within the LVZ were arbitrary and fixed at 450 km, 350 km, and
5% with respect to PREM, respectively (Fig. 10b). The result showed that the observed amplitudes were better explained (with a larger VRamp) when the LVZ was located on the western side of the ULVZ, as compared with when it was located just above the ULVZ; although, the testing of more models with different LVZ parameters is necessary to confirm this result. 7. Discussion One of the motivations of this study was to check whether the observed amplitudes were sensitive to the proposed LVZ (TFT2011), located adjacent to the ULVZ, or to the conduit of a slow region extending into the lower mantle above the ULVZ in the tomographic model (i.e., the broad dome-like plume of French and Romanowicz, 2015). So far, the results suggest that the amplitude anomalies were not produced by the conduit structure above the ULVZ (Fig. 10). Instead, the cause of the observed amplitude anomalies appears to be located to the west of the ULVZ (i.e., the earthquake side), which is in agreement with the location of the LVZ proposed by TFT2011. The height of the LVZ in Model 4 was 100 km, much lower than the one previously proposed (350 km obtained from the 2D study in TFT2011). However, since the observed amplitude anomaly at a short distances (<95°) remained unexplained in our proposed model, this height does not seem to be well constrained. Some models examined in this study explained the amplitude reduction observed at large distances (>100°); however, large

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Fig. 8. Distribution of amplitude and travel time anomalies obtained for Model 4. (a) Amplitude anomalies at 0.027–0.11 Hz. (b) Travel time anomalies at 0.027–0.11 Hz. (c) Amplitude anomalies at 0.01–0.03 Hz. (d) Travel time anomalies at 0.01–0.03 Hz.

discrepancies remained between the observed and synthetic waveforms at shorter distances (95°) for the higher frequency measurements (Figs. 2 and 8a). For the entire azimuthal range at the short distances, the observed amplitudes were much smaller than the synthetic amplitudes, which may indicate the existence of a thick slow region covering this wide azimuthal range. In Fig. S6, we show predicted amplitude anomalies for one such model. The structure has no azimuthal dependence and has a 350-km thick low Vs region at a distance of 19–50° from the earthquake. The Vs profile within this region is shown in Fig. 10b. This model explains the low amplitude observed around a distance of 95°; however, the amplitudes at larger distances were much larger than those in the observed data, therefore some modifications to the model are required. A larger set of models needs to be examined to further investigate these observations. The anomalously low S/Sdiff amplitudes observed and the emergence of postcursors (Fig. 2S) were both consequences of an anoma-

lous structure in the northern Pacific LLSVP; however, they likely have different causes. The results of beam forming analysis showed that the postcusors arrived from the south (To and Capdeville, 2011; CR2012), with respect to the station-to-earthquake backazimuth. On the other hand, the S/Sdiff amplitude was lowest at stations located to the north (azimuth 55°) of those where the postcursors were observed (azimuth 60°). If the missing energy creating the anomalously low S/Sdiff amplitude zone arrived as a postcursor from the south, the wave energy should have laterally changed its propagation direction twice, once from north to south, and then from south to north, and the amplitude of the main phase and postcursor should have shared an inverse relationship. In other words, a structural model with a large reduction in main phase amplitude should show postcursors with large amplitudes. So far, such a phenomenon has not been seen in synthetic waveforms. As shown in Fig. 7, the LVZ of our modified model reduced the amplitude of both the main phase and the postcursors.

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were mainly generated by the wave deflection effect of the elastic structure. We showed the importance of using full 3D synthetic waveforms; although, the high numerical cost of calculating the full 3D synthetic waveforms of up to 0.125 Hz prevented us from testing more number of models. On the other hand, we also showed that amplitude anomalies are visible at frequencies as low as 0.05 Hz (e.g., Fig. 2b). A future direction of study will be to lower frequency range for the synthetic waveform calculations and then to examine a larger number of models created by systematically changing the parameters that define the LVZ (e.g., horizontal location, height, Vs profile within the LVZ, and shape). 8. Conclusions

Fig. 9. Distribution of the synthetic amplitude anomalies obtained for the cylindrical LVZ structure of Model 4 with PREM as the background model.

In addition to the wave deflection effect of the 3D elastic structure, the anelasticity of the sampled structure should also affect waveform amplitudes. Some recent studies have suggested that shear attenuation of the lowermost mantle is lower than that of PREM (i.e., higher in terms of the quality factor Ql; Hwang and Ritsema, 2011; Durand et al., 2013); therefore, globally investigated Ql does not explain the observed low amplitude anomalies in this study. Although a localised low Ql structure may exist and be partly responsible for the observed low amplitudes, it should not be a primary cause because a small shift of the event location drastically changed the spatial pattern of the observed amplitude anomaly (Fig. 6). The high sensitivity of the anomalies to the direction of the path validated our assumption that they

Along with previously reported anomalous travel times and postcursors, the S/Sdiff waveforms recorded in North America for earthquakes in Papua New Guinea show anomalously low amplitudes at relatively high frequencies ( J 0.03 Hz). In this study, we showed that these amplitude anomalies were not generated by errors in the focal mechanism used to calculate predicted amplitudes. A comparison of the spatial distributions of amplitude anomalies for two different but closely located earthquakes indicated that the cause of the amplitude anomalies is most likely located in the lower mantle, rather than directly beneath the stations or earthquakes. Furthermore, the observed anomalies were found to be highly sensitive to slight shifts in the sampling path, which indicated that the anomalies were mainly generated by wave deflection effects of an elastic 3D structure, rather than from a zone of increased anelastic attenuation. We also investigated whether the amplitude anomalies could be explained by the wave deflection effect of existing Vs models and some modified models. Full 3D synthetic waveforms were calculated for about 30 structural models and then compared with observed waveforms. The results indicated that existing models can explain the trend of the observed amplitude anomalies, especially at large distances; however, the sizes of anomalies were under-predicted. One of our best models (Model 4) had a low

Fig. 10. (a) Eleven structural models with a LVZ placed on the CMB with the ULVZ (black dotted line) of CR2012 in the background. The solid circles show the tested locations of the LVZ, with colours indicating the amplitude variance reduction of each model. The amplitude anomalies for the model with the largest variance reduction (VRtt = 74.5%; VRamp = 27.9%), whose LVZ location is indicated by a thick light blue line, are plotted at station locations. For the region where the ULVZ and LVZ overlap, the Vs reduction of the ULVZ is used. (b) The Vs profile within the LVZ. The Vs lineally decreases from 350 to 280 km above the CMB. At the CMB, Vs reduction is 5% with respect to PREM.

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velocity zone, with less Vs reduction than the ULVZ, placed on the southwest side of the ULVZ of CR2012, with a tomographic model in the background. The location of the LVZ mostly overlapped with the northern Pacific LLSVP and the results indicated that the LLSVP has a larger Vs gradient at the boundary than that provided in the tomographic model. Wave deflection can occur at the boundary toward the interior of the low shear velocity province. While the model explained the low amplitudes at southern stations with distances of larger than 100°, the anomalously low amplitudes observed at distances of around 95° from the earthquake remained unexplained. Compared with travel time anomalies, amplitude anomalies should be sensitive to the spatial gradient of the 3D structure; therefore, they represent useful measurements for constraining the 3D structure. The future direction of this study is to examine a larger number of structural models, constructed by systematically changing parameters that define the LVZ (e.g., shape, location, height, and Vs profile within the LVZ). Such studies would help to better constrain the causes of the observed amplitudes, and in particular spatial gradients in the lower mantle Vs structure. Acknowledgments AT was funded by a Research Fellowship of the Japan Society for the Promotion of Science (JSPS) for Grants-in-Aid for Young Scientists (B: 23740345). BR acknowledges support from NSF grant EAR-1417229. AT thanks Nozomu Takeuchi and Alexander Nichols for fruitful discussions. We thank two anonymous reviewers and Vernon Cormier (Editor) for detailed and constructive comments. We also thank the Incorporated Research Institution for Seismology for providing waveform data. Synthetic waveform calculations were performed using the SGI ICE X of JAMSTEC SC System. Appendix A. Supplementary data Supplementary data associated with this article can be found, in the online version, at http://dx.doi.org/10.1016/j.pepi.2016.04.001. References Capdeville, Y., To, A., Romanowicz, B., 2003. Coupling spectral elements and modes in a spherical earth: an extension to the ‘sandwich’ case. Geophys. J. Int. 154, 44–57. Cottaar, S., Romanowicz, B., 2012. An unusually large ULVZ at the base of the mantle near Hawaii. Earth Planet. Sci. Lett. 355–356, 213–222. Durand, S., Matas, J., Ford, S., Ricard, Y., Romanowicz, B., Montagner, J.-P., 2013. Insights from ScS–S measurements on deep mantle attenuation. Earth Planet. Sci. Lett. 374, 101–110. Dziewonski, A.M., Anderson, D.L., 1981. Preliminary reference Earth model. Phys. Earth Planet. Int. 25, 297–356. Ekström, G., Nettles, M., Dziewonski, A.M., 2012. The global CMT project 2004– 2010: centroid-moment tensors for 13,017 earthquakes. Phys. Earth Planet. Inter. 200–201, 1–9. French, S.W., Romanowicz, B.A., 2014. Whole-mantle radially anisotropic shear velocity structure from spectral-element waveform tomography. Geophys. J. Int. 199, 1303–1327. French, S.W., Romanowicz, B., 2015. Broad plumes rooted at the base of the earth’s mantle beneath major hotspots. Nature 525, 95–99. Garnero, E.G., McNamara, A.K., 2008. Structure and dynamics of Earth’s lower mantle. Science 2008, 626–628. Garnero, E.J., Revenaugh, J., Williams, Q., Lay, T., Kellogg, L.H., 1998. Ultralow velocity zone at the core-mantle boundary. In: Gurnis, M., Wysession, M.E.,

Knittle, E., Buffett, B.A. (Eds.), The Core-Mantle Boundary Region, Am. Geophys. Un. http://dx.doi.org/10.1029/GD028p0319. He, Y., Wen, L., 2009. Structural features and shear-velocity structure of the ‘‘Pacific Anomaly”. J. Geophys. Res. 114, B02309. http://dx.doi.org/10.1029/ 2008JB005814. Houser, C., Masters, G., Shearer, P., Laske, G., 2008. Shear and compressional velocity models of the mantle from cluster analysis of long-period waveforms. Geophys. J. Int. 174, 195–212. Hwang, Y.H., Ritsema, J., 2011. Radial Ql structure of the lower mantle from teleseismic body-wave spectra. Earth Planet. Sci. Lett. 303, 369–375. Idehara, K., 2011. Structural heterogeneity of an ultra-low-velocity zone beneath the Philippine Islands: implications for core–mantle chemical interactions induced by massive partial melting at the bottom of the mantle. Phys. Earth Planet. Inter. 184, 80–90. http://dx.doi.org/10.1016/j.pepi.2010.10.014. Kennett, B.L.N., Gudmundsson, O., 1996. Ellipticity corrections for seismic phases. Geophys. J. Int. 127, 40–48. Kustowski, B., Ekstrom, G., Dziewonski, A., 2008. Anisotropic shear-wave velocity structure of the Earth’s mantle: a global model. J. Geophys. Res. 113, 303–318. Lay, T., Garnero, E.J., 2011. Deep mantle seismic modelling and imaging. Annu. Rev. Earth Planet. Sci. 39, 91–123. Lekic, V., Cottaar, S., Dziewonski, A., Romanowicz, B., 2012. Cluster analysis of global lower mantle tomography: a new class of structure and implications for chemical heterogeneity, Earth Planet. Sci. Lett. 357, 68–77. Ni, S., Tan, E., Gurnis, M., Helmberger, D., 2002. Sharp sides to the African superplume. Science 296 (5574), 1850–1853. Pachhai, S., Dettmer, J., Tkalcˇic´, H., 2015. Ultra-low velocity zones beneath the Philippine and Tasman Seas revealed by a trans-dimensional Bayesian waveform inversion. Geophys. J. Int. 203, 1302–1318. Pulliam, J., Grand, S., Sansom, J., 2008. Seismic Investigation of Edge Driven Convection Associated with the Rio Grande Rift. International Federation of Digital Seismograph Networks. Other/Seismic Network, 10.7914/SN/XR_2008, 10.7914/SN/XR_2008. Ritsema, J., Deuss, A., van Heijst, H., Woodhouse, J., 2011. S40RTS: a degree-40 shear-velocity model for the mantle from new Rayleigh wave dispersion, teleseismic traveltime and normal-mode splitting function measurements. Geophys. J. Int. 184, 1223–1236. Romanowicz, B., 1987. Multiplet-multiplet coupling due to lateral heterogeneity: asymptotic effects on the amplitude and frequency of the Earth’s normal modes. Geophys. J. Roy. Astron. Soc. 90, 75–100. Romanowicz, B., Mitchell, B., 2015. Deep Earth structure: Q of the Earth from crust to core. In: Schubert, G. (Ed.), Treatise on Geophysics. Elsevier, pp. 789–827. Rondenay, S., Cormier, V.F., Van Ark, E.M., 2010. SKS and SPdKS sensitivity to twodimensional ultralow-velocity zones. J. Geophys. Res. 115. http://dx.doi.org/ 10.1029/2009JB006733. Takeuchi, N., 2007. Whole mantle SH velocity model constrained by waveform inversion based on three-dimensional Born kernels. Geophys. J. Int. 169, 1153– 1163. Takeuchi, N., Geller, R.J., Cummins, P.R., 1996. Highly accurate P-SV complete synthetic seismograms using modified DSM operators. Geophys. Res. Lett. 23, 1175–1178. Thorne, M.S., Zhang, Y., Ritsema, J., 2013a. Evaluation of 1-D and 3-D seismic models of the Pacific lower mantle with S, SKS, and SKKS traveltimes and amplitudes. J. Geophys. Res. Solid Earth 118. http://dx.doi.org/10.1002/jgrb.50054. Thorne, M.S., Garnero, E.J., Jahnke, G., Igel, H., McNamara, A., 2013b. Mega ultra low velocity zone and mantle flow. Earth Planet. Sci. Lett. 364, 59–67. To, A., Capdeville, Y., 2011. Constraints on the 3D shape of the ultra low shear velocity zone at the base of the mantle beneath the Central Pacific. AGU 2011 Fall Meeting San Francisco, December 5–9, DI43A–2082. To, A., Romanowicz, B., 2009. Finite frequency effects on global S diffracted traveltimes. Geophys. J. Int. 179, 1645–1657. To, A., Romanowicz, B., Capdeville, Y., Takeuchi, N., 2005. 3D effects of sharp boundaries at the borders of the African and Pacific superplumes: observation and modelling. Earth Planet. Sci. Lett. 233 (1–2), 1447–1460. To, A., Fukao, Y., Tsuboi, S., 2011. Evidence for a thick and localized ultra low shear velocity zone at the base of the mantle beneath the Central Pacific. Phys. Earth Planet. Int. 184, 119–133. Woodhouse, J.H., Wong, Y.K., 1986. Amplitude, phase and path anomalies of mantle waves. Geophys. J. Roy. Astron. Soc. 87, 753–773. Zhang, Y., Ritsema, J., Thorne, M., 2009. Modelling the ratios of SKKS and SKS amplitudes with ultra-low velocity zones at the core-mantle boundary. Geophys. Res. Lett. 36. http://dx.doi.org/10.1029/2009GL040030. Zhao, C., Garnero, E.J., McNamara, A.K., Schmerr, N., Carlson, R.W., 2015. Seismic evidence for a chemically distinct thermochemical reservoir in Earth’s deep mantle beneath Hawaii. Earth Planet. Sci. Lett. 426, 143–153.