2D spatial distribution of reflection intensity on the upper surface of the Philippine Sea plate off the Boso Peninsula, Japan

Journal Pre-proof 2D spatial distribution of reflection intensity on the upper surface of the Philippine Sea plate off the Boso Peninsula, Japan

Akihiro Kono, Toshinori Sato, Masanao Shinohara, Kimihiro Mochizuki, Tomoaki Yamada, Kenji Uehira, Takashi Shinbo, Yuya Machida, Ryota Hino, Ryousuke Azuma PII:

S0040-1951(19)30321-X

DOI:

https://doi.org/10.1016/j.tecto.2019.228206

Reference:

TECTO 228206

To appear in:

Tectonophysics

Received date:

29 January 2019

Revised date:

14 September 2019

Accepted date:

15 September 2019

Please cite this article as: A. Kono, T. Sato, M. Shinohara, et al., 2D spatial distribution of reflection intensity on the upper surface of the Philippine Sea plate off the Boso Peninsula, Japan, Tectonophysics(2019), https://doi.org/10.1016/j.tecto.2019.228206

This is a PDF file of an article that has undergone enhancements after acceptance, such as the addition of a cover page and metadata, and formatting for readability, but it is not yet the definitive version of record. This version will undergo additional copyediting, typesetting and review before it is published in its final form, but we are providing this version to give early visibility of the article. Please note that, during the production process, errors may be discovered which could affect the content, and all legal disclaimers that apply to the journal pertain.

© 2019 Published by Elsevier.

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2D spatial distribution of reflection intensity on the upper surface of the Philippine Sea plate off the Boso Peninsula, Japan Akihiro Konoa, *,+, Toshinori Satoa, Masanao Shinoharab, Kimihiro Mochizukib, Tomoaki Yamadab, Kenji Uehirac, Takashi Shinbod, Yuya Machidad, Ryota Hinoe, Ryousuke Azumae a

Graduate School of Science, Chiba University, 1-33, Yayoicho Inage-ku, Chiba 263-8522, Japan

*Corresponging author +Now at INPEX Co.

Sato:

[email protected]

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Kono:

[email protected]

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b

Earthquake Research Institute, University of Tokyo, 1-1-1 Yayoi, Bunkyo-ku, Tokyo 113-0032, Japan

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National Research Institute for Earth Science and Disaster Prevention, 3-1 Tennodai, Tsukuba, Ibaraki 305-0006, Japan

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Japan Agency for Marine-Earth Science and Technology, 25-3173 Showamachi, Kanazawa-ku, Yokohama 236-0001, Japan

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Graduate School of Science, Tohoku University, 6-6 Aramaki Aza Aoba, Sendai, Miyagi 980-8578, Japan

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Abstract

The region off the Boso Peninsula, Japan, is a tectonically complex area where the Pacific plate is subducting beneath both the landward plate and the Philippine Sea plate (PHS)

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from the Japan trench and the Izu-Bonin trench as the PHS is subducting under the landward plate from the Sagami trough. It is important to better determine the structure of this region to deepen our understanding of its seismicity. Previous seismic reflection studies

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have shown that reflections from the upper surface of the PHS vary with depth, being stronger in the main slip area of the slow slip events beneath the Boso Peninsula (Boso SSEs). However, the spatial relationship between the reflective area and the SSEs is poorly constrained. This study mapped the distribution of the reflective area using data recorded by ocean bottom seismometers during an active-source seismic experiment. We constructed a 3D P-wave velocity structure by using traveltimes of first arrivals from 18 ocean bottom seismometer records. We also adapted the traveltime mapping method to reflection traveltimes, projecting them to the depth–distance domain, to map the 2D distribution of strong reflections from the top of the PHS. These reflections were concentrated in two areas, one near the main slip area of the Boso SSEs and the other about 60 km to the east. In the first area, the absence of strong velocity contrasts near the top of the PHS suggests that the reflections were generated by a thin low-velocity layer. In contrast, the structure of the second area has a convex shape of high velocity with a high velocity gradient near the top of the PHS. This structure may represent boninitic material of the outer-arc high, partially

serpentinized peridotite, or gabbro displaced by intraoceanic reverse faults.

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Highlights 3D traveltime mapping shows two strong reflection areas from the top of the PHS. One is near the main slip area of the Boso SSEs and the other is about 60 km to the east. The reflections from the former area are generated by a thin low-velocity layer. The latter area contains a high-velocity structure (HVS) near the top of the PHS. This HVS may represent boninite or partially serpentinized peridotite or gabbro. Keywords Philippine Sea plate, Traveltime mapping, Ocean bottom seismometers (OBSs), 3D P-wave

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velocity structure, Boso slow slip events (SSEs), Izu-Bonin arc

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1. Introduction

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The region off the Boso Peninsula, central Japan, has a complex tectonic setting in which two oceanic plates are subducting beneath the landward plate on which the Japanese islands sit. The Pacific plate (PAC) is subducting west-northwestward under both the Philippine Sea plate (PHS) and the landward plate at the Japan trench and the

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Izu-Bonin trench, and the PHS is itself subducting northwestward under the landward

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plate at the Sagami trough (Fig. 1). This tectonic setting has led to diverse seismic

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events in the region. Major interplate earthquakes have repeatedly occurred near the

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Sagami trough along the plate boundary between the PHS and the landward plate (e.g.,

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the 1703 Genroku Kanto earthquake and the 1923 Taisho Kanto earthquake), which

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have caused great damage and many casualties in the Tokyo metropolitan area.

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This plate boundary has long been the subject of scientific investigations using a variety of techniques. Global Navigational Satellite System stations have yielded geodetic evidence of slow slip events (SSEs) off the Boso Peninsula (e.g., Ozawa et al., 2003, 2007; Sagiya, 2004; Ozawa, 2014; Fukuda et al., 2014). These Boso SSEs, which last 10–20 days and occur at intervals of 2 to 7 years, have been documented since the 1990s. They are located along the upper surface of the PHS (UPHS), the boundary between the PHS and the landward plate (Fig. 1b). To better interpret these seismic and aseismic events, we need to delineate the tectonic structure of the Boso Peninsula region. 4

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Previous studies have conducted active-source and passive-source seismic experiments for that purpose. Passive seismic studies have estimated the depth of the UPHS in the region (e.g., Ishida, 1992; Hori, 2006; Kimura et al., 2006; Hirose et al., 2008; Uchida et al., 2009,

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2010). However, the UPHS is poorly defined in the region off the Boso Peninsula

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because seismicity is relatively sparse there and because most seismic data are recorded

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at distant land-based stations. Ito et al. (2017a, 2017b) collected data with ocean bottom

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seismometers (OBSs) off the east and southeast coast of the Boso Peninsula. Although

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they succeeded in modeling the 3D seismic velocity structure in the region, the

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resolution at shallower depths (~15 km) southeast of the Boso Peninsula was poor given

this area.

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the low seismicity. Therefore, passive seismic tomography has been of limited value in

Other studies have used active seismic experiments. Sato et al. (2005) estimated the depth of the UPHS beneath the Kanto region, including the Boso Peninsula, on the basis of a 2D seismic reflection survey. Tsumura et al. (2009) imaged the UPHS down to 18 km depth under the southernmost Boso Peninsula. Kimura et al. (2009) imaged the UPHS to a depth of 20 km using data obtained by 2D seismic reflection experiments off the Boso Peninsula. Nakahigashi et al. (2012) conducted a 2D refraction survey and 5

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determined the 2D P-wave velocity structure beneath the offshore region between the southern Tohoku region and the Boso Peninsula, imaging the depths of the PHS and PAC. Kono et al. (2017) modeled the 2D velocity structure, including the depth of the UPHS, under the region off the Boso Peninsula and integrated it with previously

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published results to draw depth contours of the UPHS.

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Several studies have analyzed the intensity of seismic reflections from the UPHS. Sato

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et al. (2005) found that the reflections from the UPHS beneath Tokyo Bay ranged in

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amplitude from strong in the north to weak in the south. Kimura et al. (2009) also found

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that seismic reflections from the UPHS off the Boso Peninsula varied with depth near

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the Sagami trough, where trough filling sediments and basaltic rocks are juxtaposed at

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the UPHS. Reflections were strong down to 10 km depth, weaker from 10 to 14 km depth, and strong below 14 km (Fig. 1b). They noted that the area of weak reflections corresponds to the source region of great interplate earthquakes and the area of strong reflections corresponds to the main slip area of the Boso SSEs. Kono et al. (2017) also reported strong reflections near the Sagami trough and the main slip area of the Boso SSEs, but observed no reflections between the two areas, and eastern area from about 141 E. Variations in reflection intensity have also been reported elsewhere on the plate 6

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boundary. In an active-source experiment using OBSs in the forearc region of the Japan trench, off the Tohoku region, Mochizuki et al. (2005) observed strong reflections from the boundary between the landward plate and the PAC, where seismicity is relatively quiet. They interpreted the highly reflective area as a thin low-velocity layer atop the

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PAC, possibly composed of material with a low frictional coefficient with the potential to

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host aseismic slip, such as hydrous minerals, aqueous fluids, or serpentinite. Although

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no SSEs have been detected in this area, the inverse relation between reflection intensity

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and seismicity implies that the distribution of reflection intensity along a plate boundary

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can help distinguish aseismic and seismic slip regions.

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No previous studies have presented a precise 3D seismic velocity structure for depths

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shallower than ~15 km off the Boso Peninsula, as the strong reflections from the shallow UPHS were limited to 2D profiles (Kimura et al., 2009; Kono et al., 2017). Determining whether the slip area of the Boso SSEs corresponds to the reflective area on the UPHS requires knowledge of the 2D spatial distribution of reflectivity on the UPHS. A 3D seismic velocity structure is also important for determining the nature of the velocity contrasts that give rise to reflections. In this study, we report a 3D seismic velocity structure and the 2D distribution of reflection intensity on the UPHS off the Boso Peninsula, and we discuss their relation to 7

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the Boso SSEs and materials in the PHS plate.

2. Data acquisition and analysis 2.1. Data acquisition

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From 30 July to 4 August 2009, the Japan Agency for Marine-Earth Science and

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Technology (JAMSTEC) ship R/V Hakuho-maru conducted an active seismic survey off

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the southeastern coast of the Boso Peninsula in which airguns were fired along four

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survey lines and reflection data were collected by an array of OBSs (Fig. 2).

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In this study, we used 29 OBSs, equipped with three-component velocity sensors that

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were packed in spheres of titanium. The OBS sensors had a natural frequency of 1 Hz

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and a 200 Hz sampling rate. The OBSs were deployed in July 2009 by the R/V Hakuho

maru and retrieved in October 2010 by the Shincho-maru of Shin-Nihon-Kaiji Co. (now Fukada Salvage Co.). All but two of the OBSs yielded useful data (Fig. 1b). The active source was an array of four Bolt 1500 LL-guns with a total volume of 6000 cubic inches (98.3 L). These guns fired 3024 times along line M, 246 times along line S1, 467 times along line S2, and 379 times along line S5 (Fig. 2). Shots were fired at intervals of about 1 min to provide a shot interval of about 100 m along the survey lines.

2.2. 3D velocity structure analysis

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We established a domain for the 3D model extending 150 km parallel to our seismic lines (axis X), 90 km perpendicular to those lines (axis Y), and 30 km in depth (axis Z) (Fig. 2). We used 2099 of 3024 shots along line M and all shots along lines S1, S2, and S5.

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We prepared the traveltime data set by picking first-arrival times in the OBS records.

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First arrivals were clearly recognized for the most part along each survey line (Fig. 3,

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top). We also picked first arrivals in records from OBSs outside the survey lines.

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Although these records were from offsets greater than 10 km, first arrivals could be

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recognized in most of them (Fig. 3, bottom). Consequently, we obtained 30,544

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traveltimes with a picking error of ±50 ms.

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To build an initial 3D velocity structure, we constructed four 2D velocity profiles across our model domain at 13, 33, 52, and 72 km on axis Y (Fig. 4). We adopted the 2D velocity model determined by Kono et al. (2017) for our profile at Y = 33 km, including the geometry of the UPHS. The other three velocity structures were built on the basis of this velocity structure. We also took into account the velocity structures estimated by Kimura et al. (2009) and Nakahigashi et al. (2012). The 2D velocity structure of Nakahigashi et al. (2012) extended under the forearc region of the Japan trench, including off the Boso Peninsula. This structure includes the upper and lower layers of 9

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the island arc crust above the UPHS in the northern part of our survey area, with estimated velocities of 5.4–5.9 km/s and 6.1–6.5 km/s, respectively. Extending that model, we inserted upper island arc crust (8–10 km thick) just above the subducting PHS in our profile at Y = 52 km and upper and lower crusts (8–9 and 3–9 km thick,

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respectively) in our profile at Y = 72 km, with the same velocities assigned by

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Nakahigashi et al. (2012) to both crusts (Fig. 4b). We built an initial 3D P-wave velocity

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model by linear interpolation among these four 2D models.

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After constructing the initial model, we applied the First Arrival Seismic Tomography

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(FAST) program (Zelt and Barton, 1998) to the first-arrival time data and the initial

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model. The 150 × 90 × 30 km model for our FAST application was composed of 301 ×

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181 × 61 nodes for the forward model and 150 × 90 × 30 nodes for the inversion. The result, after 10 iterations, was our final 3D P-wave velocity model. To confirm the reliability of the final model, we conducted a checkerboard resolution test. We added velocity perturbations of ±10% to alternating 20 × 20 × 6 km blocks of the final model (Fig. 5), then calculated the synthetic traveltimes of the perturbed model and used them as input data for the FAST inversion. We used the difference in velocities between the resulting model and the original observed data to assess whether the added perturbations were well recovered. A well-recovered checkerboard is an indication that 10

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the velocity model is reliable.

2.3. 3D traveltime mapping method for reflection phases We detected later phases in some of the OBS records (purple lines in Fig. 3). From these, we selected phases with curves similar to the parabolic shape of seismic reflections

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or with apparent velocities significantly greater than those of refracted first arrivals. The

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picked phases also had significant amplitudes, meaning that they remained appreciable

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when normalized against those of the first arrivals. We made a total of 2099 picks of

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reflection phases, which were recorded at 15 OBSs.

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We applied the traveltime mapping method developed by Fujie et al. (2006) to the

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selected phases to estimate the locations of reflectors. This method, originally developed

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for 2D models, has not previously been applied to a 3D space. The traveltime mapping method assumes a paired source (S) and receiver (R) in the model and a traveltime of a reflection phase (T0) observed in an OBS record. The diffraction stacking method considers the phase as a superposition of diffractions at any point P where the total traveltime from S to R via point P is equal to T0. The locus of P, which is named the diffraction surface, can be determined by summing the two traveltime fields from S and R. The reflection point of the observed phase should be somewhere on the diffraction surface. The actual reflection point is likely located near the midpoint between the source (𝑥𝑠 , 𝑦𝑠 , 𝑧𝑠 ) and receiver (𝑥𝑟 , 𝑦𝑟 , 𝑧𝑟 ) in the general structure. The likelihood is represented by weight functions of the

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Gaussian distribution with mean 𝑥𝑚 = (𝑥𝑠 + 𝑥𝑟 )/2 and 𝑦𝑚 = (𝑦𝑠 + 𝑦𝑟 )/2 . The imaging condition 𝐴(𝑥, 𝑦, 𝑧) can be calculated by 𝐴(𝑥, 𝑦, 𝑧) =

1 −∆𝑇 2 /2𝜏2 1 −𝑑2 /2𝜃2 𝑒 𝑒 √2𝜋𝜃 √2𝜋𝜏

(1),

where τ is the uncertainty of the picked traveltime; ∆𝑇 = 𝑇0 − 𝑡(𝑥, 𝑦, 𝑧): 𝑡(𝑥, 𝑦, 𝑧) is the

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traveltime

at

( 𝑥, 𝑦, 𝑧 ); 𝑑 = √(𝑥𝑚 − 𝑥)2 + (𝑦𝑚 − 𝑦)2 ;

and

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𝜃 = 𝛼√(𝑥𝑠 − 𝑥𝑟 )2 + (𝑦𝑠 − 𝑦𝑟 )2 is the aperture of the ordinary migration method,

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which adjusts the sharpness of the image. According to Fujie et al. (2006), α ranges

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from 0.1 to 0.3 for most cases, the value depending on the crustal structure. In this

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study, as we varied the value of α between 0.1 and 0.3, we could not recognize significant difference from the result. Hence, we applied the value of α to be 0.3 in

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whole area.

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Reflectors can be imaged by stacking the Gaussian weighted diffraction surfaces of each picked reflection traveltime. The resulting image consists of multiple gray

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clouds, their fuzzy outlines representing the uncertainty of the reflectors’ locations. The late arrivals of the reflections are usually more difficult to pick precisely than first arrivals owing to the presence of competing phases. In this study, the later phases were picked with an uncertainty of 100 ms, double the error of the first-arrival picks.

3. Results 3.1. 3D velocity model The root-mean-square traveltime residuals between synthetic and observed 12

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first-arrival times for the initial 3D model ranged from 347 ms, with a 𝜒 2 value of 48.1. After 10 iterations of 3D inversions, which yielded the final velocity model (Fig. 6), the residuals decreased to approximately 50 ms, with a 𝜒 2 value of almost 1. As shown in Fig. 3, we can see good agreement with the observed and synthetic traveltimes.

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Kono et al. (2017) estimated the depth of the UPHS off the Boso Peninsula by

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integrating estimates from previous passive and active seismic studies. We

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superimposed their UPHS model on our final 3D velocity structure (Fig. 6b, c) and

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found that the UPHS corresponds to an area with velocities of about 5–6 km/s at deeper

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than 10 km (e.g., the UPHS at about 10 km depth along X = 60 km section in Fig. 6b).

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This indicates that we could obtain 2D configuration of the UPHS from our 3D

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structure model. The checkerboard resolution test (Fig. 7) recovered the perturbations well over a wide area at 7 and 10 km depth, and moderate at 15 and poor at 19 km depth, although recovery was poor at the southwest and northeast ends. In vertical cross sections, the perturbations appeared to be recognizable at about 15 km depth. These results indicate that the velocity model above the UPHS is reliable enough to discuss the reflections from the UPHS. We extended the UPHS by Kono et al. (2017) to the eastern area (X > 75 km, Fig. 6) on the basis of the velocity of the UPHS (5-6 km/s) and the results of resolution test.

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We also extended the UPHS to the shallower part (depth < 10 km) with the structures estimated by Kimura et al. (2009) and Nakahigashi et al. (2012) (Fig. 6).

3.2. 3D traveltime mapping We mapped the 2099 traveltimes of reflection events into the 3D space of our model

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(Fig. 8). The traveltime mapping method projects each picked reflection traveltime into

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the depth space, where it is represented as a gray cloud. We took a range of gray clouds

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as about one standard deviation of 𝐴(𝑥, 𝑦, 𝑧) in eq (1) (about ±2 km in horizontal,

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±1.5 km in vertical). The reflectors lie somewhere in these clouds, and their presence is

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more strongly indicated where multiple clouds overlap each other.

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Seen in plan view, the result of our traveltime mapping shows many reflectors spread

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across the model (Fig. 8a). Vertical sections suggest that the reflectors lie at various depths (Fig. 8b, c). By plotting the depth of the UPHS on these vertical sections, we can recognize that many reflectors are within the overriding plate, where they may represent boundaries within sedimentary layers. There are also many reflectors close to the UPHS that appear to be concentrated in the western (X = 20-50 km) and eastern (X = 90-95 km) parts of the model (Fig. 8b, c). Reflectors below the UPHS may represent a boundary within the slab, such as the boundary between the upper and lower oceanic crust or the Moho. To focus on the mapping result near the UPHS, Fig. 9a shows the 14

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traveltime mapping results within ±2.5 km of the UPHS in consideration of uncertainties of traveltime mapping (about ±1.5 km) and location of the UPHS (about ±2 km). Fig. 9b

shows the range within which a reflection from the UPHS could be observed at one of the OBS stations. We show areas only where we estimated the UPHS location. This

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observable range is determined based on geometry of survey lines and OBS locations.

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Since detection ability of the reflection phase from the UPHS may be low in some parts

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due to the overlapping with other types of phases, observable range of reflection from

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the UPHS may be limited. The results indicate that strong reflections from the UPHS

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are mainly concentrated in the western and the eastern parts of the model.

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The presence of reflections suggests that a relatively strong velocity contrast, negative

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or positive, exists along or near the UPHS. To analyze the cause of the reflections, we compare the traveltime mapping results with our seismic velocity structure and previous studies in the following section.

4. Discussion 4.1. 3D velocity structure Kimura et al. (2009) estimated that the UPHS velocity ranges from 4.8 to 5.2 km/s in their velocity analysis at around 10 km depth. At that depth, the velocity structures of 15

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Nakahigashi et al. (2012) and Kono et al. (2017) estimated the velocity of the UPHS as 4.6 km/s and 5.0 km/s, respectively. In sum, estimates of the P-wave velocity of the UPHS range from 4.6 to 5.2 km/s at around 10 km depth, and the velocities around the UPHS in our velocity model at that depth (Fig. 6) are compatible with velocity ranges

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estimated by previous studies. Our model revealed spatially continuous velocity

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structure crossing 2D survey lines of the previous studies.

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4.2. Western reflective area

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In Fig. 10a, we superimposed the slip area of the 2013-2014 Boso SSE (Sato et al.

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2017) and the reflection intensity of the UPHS along the 2D profile of Kimura et al.

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(2009) on our 3D traveltime mapping results within ±2.5 km depth from the UPHS.

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The figure indicates that the western reflective area on the UPHS lies in and around the slip area of the Boso SSE. However, the distribution of reflectors in that area extends farther south than the SSE slip area estimated by Sato et al. (2017). They estimated the slip distribution by supplementing land-based data with data from two ocean bottom pressure gauges (OBPs) placed off the Boso Peninsula (Fig. 10a), one (P1) of which detected transient vertical displacement of the seafloor during the 2013–2014 Boso SSE. Since they used only two OBPs located south-east edge of the slip area and no OBPs were located at the southern part, the slip distribution of the SSE at the southern part 16

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may remain uncertain. The velocity structure around the UPHS does not include a strong velocity contrast at around 10 km deep area (Fig. 10b, c). Although we can see a slightly upward high velocity (about 7.0 km/s) area with high velocity gradient at about X = 45-80 km, and

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about 16 km depth (Fig. 10b), this area is separated from the UPHS. We assume that

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the density contrast may be similar to the velocity contrast. Because reflections are

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generated by impedance (Vp times density) contrasts, it may be that the observed

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reflections in the western reflective area were generated by a small-scale structure on or

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near the UPHS that our study cannot resolve.

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A structure that generates a reflection can have a velocity that is either higher or lower

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than its surroundings. If the reflections were generated by a thin high-velocity layer, the velocity of that layer should be higher than that of the oceanic crust, which is 4.6–5.5 km/s off the Boso Peninsula. This would limit the likely material of the layer to gabbro, peridotite, or volcanic rocks. However, no example is known of a thin layer of such materials atop a subducting plate. Moreover, the velocity structure shows no sign of intrusions beneath the area of strong reflections. The reflections, then, are probably generated by a thin low-velocity layer atop the UPHS. Mochizuki et al. (2005) proposed that a thin low-velocity layer atop the subducting 17

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slab explains reflections from the plate boundary in an aseismic region off the Tohoku region. They suggested that the aseismic region would have aseismic slip, and a layer consisting of material with low frictional coefficient such as hydrous minerals, clay minerals, or serpentinite would account for the aseismic slip. In the region off the Boso

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Peninsula, Kono et al. (2017) similarly proposed that the reflections in the slip area of

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the Boso SSEs (light blue lines in Fig. 9a) are produced by a thin layer on the UPHS

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with a relatively low velocity of 4.0–4.5 km/s, perhaps a layer filled with trapped fluids,

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clay minerals, or buoyancy-driven serpentinite derived from the mantle wedge. If the

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southern strong reflective area around 10 km depth in this study has such a thin layer,

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this area may have potential of SSE slip.

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The distribution of reflective areas may show that the weak reflection area estimated by Kimura et al. (2009) does not extend to the southeast. If the source region of interplate earthquakes coincides with weak reflection area, this limited distribution of weak reflection area seems to be consistent with the location of the eastern edge of the source region of 1703 earthquake (Fig .1b, Sato et al., 2016) and the source region of largest aftershock of the 1923 earthquake (Kimura et al., 2009). Interplate earthquakes such as the 1703 and 1923 earthquakes have repeatedly occurred near the site of the Boso SSEs (Fig. 1) at depths of 10–30 km, similar to the 10–20 km

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depth range of the Boso SSEs (Fig. 2). It is still uncertain why these different types of events, which represent different fault slip rates, occur at similar depths (similar temperature and pressure conditions) along the plate boundary in this region. The controlling factor may be a structural feature in the subducting plate that can generate

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strong reflections and induce SSEs. A plausible explanation is the presence of materials

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with low frictional coefficient and low velocity, as suggested by Kono et al. (2017).

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4.3. Eastern reflective area

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We also detected an area of strong reflections in the eastern part of the model (Figs.

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9a, 11a). On the other hand, Kono et al. (2017) did not show the reflections under the

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Line M in the eastern part. One of reasons of this discrepancy may be that spatial

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distribution of reflections is sporadic in these studies. This is because observable range of reflection from the UPHS may be limited as mentioned Section 3.2. The velocity structure of this area (X = 85-110 km) has a distinctive convex shape of high velocity (about 6-7 km/s) with a high velocity gradient (an area of the dashed light blue

ellipse in Fig. 11b). From this structure, our result suggests a possibility that the strong reflections were generated by a large positive velocity contrast. However, there is other possibility that the reflections were generated by a negative velocity contrast with an unresolvable small-scale structure as discussed in the previous Section.

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4.4 High velocity structures We suggested that the distinctive convex shape of the high velocity structure in the eastern reflective area (X = 85-110 km) is responsible for the strong reflection. We also mentioned the slightly upward high velocity structure at about X = 45-80 km, and about

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16 km depth (Fig. 10b). We would like to discuss these structures in this section.

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Kono et al. (2017) showed a convex shape of high velocity (6-7 km/s) in these areas

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(under 120 and 160 km horizontal distances in their Fig. 4), where the resolution is

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moderate. Another previous tomographic studies may have recognized a high-velocity

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structure in the area. Nakajima and Hasegawa (2010) found a well-resolved velocity

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contrast within the PHS slab under the Boso Peninsula between a high-velocity body to

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the west and a low-velocity body to the east (their Fig. 7a to 7c). They proposed that a high-velocity body under the Boso Peninsula lies between the PHS and the PAC, extending to the upper surface of the PHS slab, and suggested that this structure has been observed in previous tomographic studies. Ito et al. (2017b) also reported a velocity contrast within the PHS slab off the Boso Peninsula. From P wave structures of depth : 15 km in Fig. 3 and Y = -60 km in Fig. 4 in Ito et al. (2017b), an area of high P-wave velocity (>7.0 km/s) below about 13-14km depth is located in our model at X = 50-70 km, and about 1 km below the UPHS (Fig. 11a), although their checkerboard

20

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resolution test suggests that their model resolution was poor at 10–20 km depth. Our model shows the slightly upward high velocity (about 7.0 km/s) structure with high velocity gradient at about X = 45-80 km, and about 16 km depth (Fig. 10b and Fig. 11b). Despite of depth differences, Ito’s high velocity structure may correspond to this area,

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because their resolutions are not good in this area.

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No previous studies have discussed the lithology or the formation mechanism of the

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high-velocity structures in the PHS slab. Here we evaluate three candidates for the

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material making up these structures: boninite, peridotite, and gabbro.

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Clues to the nature of the high-velocity structures may be found in the Izu-Bonin

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island arc, which lies in the eastern part of the PHS and is currently subducting

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northwestward off the Boso Peninsula at the Sagami trough (Fig. 1). In their 2D P-wave velocity structure along a section across the Izu-Bonin arc, Suyehiro et al. (1996) noted a convex upward high-velocity structure (at about 141.3E and 7 km/s at 7 km depth) and a serpentinite seamount near the Izu-Bonin trench. From a comparison of this velocity structure and the seismic reflection image of the Izu-Bonin arc by Taylor (1992), it appears that the convex upward structure corresponds to the outer-arc highs formed by boninitic volcanic activity during the Eocene. Suyehiro et al. (1996) placed the boundary between the convex upward structure and the serpentinite seamount almost 21

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40 km west of the trench. The convex upward structure is about 50 km wide. To project the position into the model area of this study, we should recognize the shape of the subducted PHS-PAC boundary, which is still on debate. Takahashi (2006) implied that the PHS-PAC boundary has been almost straight since 15 Ma from the time when

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Izu-Bonin arc collides with the Japan island arc around the area. On the other hand,

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some studies implied that the PHS plate deforms after subduction (e.g. Uchida et al.,

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2010), because the thin tip of the PHS may be able to deform easily. Hence, in Fig. 11a,

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we show two patterns of the area of the convex upward structure. One is followed by the

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straight PHS-PAC boundary, the other is followed by the deformed boundary. We found

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that both high-velocity structures around X = 60, 100 km almost coincides with both

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areas of the projected convex upward structure (Fig. 11a). The second candidate material for the high-velocity structures is partially serpentinized peridotite. Dehydration of the subducting PAC slab supplies large amounts of water to the mantle wedge in the Izu-Bonin arc, which promotes serpentinization of peridotite (e.g., Hyndman and Peacock, 2003). Previous studies have suggested that partially serpentinized peridotite rises from the mantle wedge to the forearc region in the Izu-Bonin arc and is extruded onto the seafloor along extensional faults (e.g., Fryer, 1996, Kamimura et al., 2002, Oakley et al., 2007, Fujioka, 2012), 22

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where it forms conical seamounts near the trench. Serpentinite seamounts are mainly composed of serpentinite mud with low velocity and density. Oakley et al. (2007) estimated a velocity of 1.6 km/s in serpentinite seamounts in the Mariana arc, and Suyehiro et al. (1996) and Kamimura et al. (2002)

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estimated velocities of ~3.5 km/s and 3.2 km/s, respectively, in a serpentinite seamount

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in the Izu-Bonin arc. Serpentinite seamounts may be scraped off the subducting PHS at

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the Sagami trough, but the partially serpentinized peridotite underlying them would

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remain in the lower plate at or near the subducting UPHS.

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Kamimura et al. (2002) produced a velocity profile in the Izu-Bonin arc near the

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profile of Suyehiro et al. (1996) and estimated the velocity of partially serpentinized

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peridotite as 6.5–6.8 km/s at around 10 km depth (almost 5 km below the seafloor), similar to the 7.0 km/s velocity of the convex upward structure of Suyehiro et al. (1996). Hyndman and Peacock (2003) compiled laboratory determinations of seismic velocities of partially serpentinized peridotite at 1 GPa, corresponding to ~35 km depth. Their evidence suggests that a P-wave velocity of 6.5–7.0 km/s corresponds to 40–55% serpentinization of peridotite (Fig. 6 of Hyndman and Peacock, 2003). This degree of serpentinization is close to the 40–60% serpentinization rate that Hyndman and Peacock reported in the uppermost mantle in the Mariana forearc. 23

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The presence of 40–55% serpentinized peridotite has not been confirmed at depths shallower than 35 km in the subducting PHS. However, Fujioka et al. (1994, 1995) recovered serpentinite from the seafloor in the Izu-Bonin forearc containing boulders of peridotite, which they interpreted as xenoliths entrained in the rising serpentinite. The

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presence of mixed peridotite and serpentinite on the seafloor implies that peridotite also

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exists at shallow depths (~5 km from the seafloor in the velocity structure of Kamimura

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et al., 2002). The high-velocity structure seen in this study is a few kilometers deeper

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than the estimated depth of the UPHS. Thus, it is plausible that the high-velocity

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structure is composed of partially serpentinized peridotite.

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One effective way to confirm the presence of serpentinite is with the Vp/Vs ratio, an

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index of elastic properties that increases with progressing serpentinization of peridotite (Christensen, 2004). Hyndman and Peacock (2003) reported that at 1 GPa pressure (~35 km depth), Vp/Vs of peridotite rises from about 1.8 to >2.1 as serpentinization proceeds from zero to 100%. In the high-velocity area modeled by Nakajima and Hasegawa (2010), Vp/Vs ranges from 1.7 to 1.9, although its location is far from the high-velocity structure imaged in this study. Ito et al. (2017b) estimated Vp/Vs values of 1.9 to 2.1 off the Boso Peninsula near the high-velocity structure of this study; however, the resolution of their velocity structure was poor in this area. 24

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The third candidate is the gabbro that makes up lower crust, which has a P-wave velocity of about 7.0 km/s by Kodaira et al. (2007), and Nakahigashi et al. (2012). Tsuji et al. (2013) reported reverse and strike-slip faults in the subducting PHS off the Kii Peninsula, southwest of the study area near the Nankai trough. Some of these faults

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extended from the seafloor to the depth of the Moho, and seismic reflection images

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showed that they raised the crust more than 1 km. Although none have been reported

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and the raise offset reported by Tsuji et al. (2013) may be insufficient to explain the

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high-velocity structure around X = 100 km, similar faults in the crust of the PHS could

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have displaced deep gabbro to positions matching the high-velocity structures if the

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PHS plate deforms after subduction.

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Further research is needed to reveal the lithology and formation mechanism of the high-velocity structures. An approach to determine the Vs structure and Vp/Vs ratio with sufficient resolution is to determine the presence of serpentinite in the structure. Althouth background seismicity in the eastern reflective area is very low and no large earthquakes or SSEs have been documented in the past, it is unknown whether slip behavior of the UPHS at the area is stable or unstable. Frictional properties of boninite and gabbro have not yet been determined at temperatures from 100 to 150 °C, which is the temperature of the UPHS at 10–20 km depth in our study area ranges (Wada and He, 25

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2017). If these frictional properties will be determined, and ocean bottom geodetic observations will be conducted to estimate slip behavior, these are helpful to estimate the lithology of the high-velocity structure around X = 100 km.

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5. Conclusions

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We conducted an active-source seismic experiment off the Boso Peninsula to

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determine the 3D P-wave velocity structure and map the 2D spatial distribution of

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reflection intensity along the UPHS. We estimated the 3D velocity structure from the

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traveltime of first arrivals in OBS records. We then picked the reflection traveltimes of

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phases with curve geometries similar to the parabolic shape of seismic reflections or with

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apparent velocities that were significantly faster than those of refracted first arrivals, and applied the traveltime mapping method to project these reflection traveltime data to the depth-distance domain. This study, the first application of the traveltime mapping method to a 3D velocity structure by traveltime mapping in 3D, enabled us to estimate the 2D spatial distribution of reflection intensity along the UPHS. Many reflections from the UPHS were observed in and around the slip area of the 2013–2014 Boso SSE and in the eastern part (X = 85-110 km) of our model region. In a comparison of the traveltime mapping results and our 3D velocity structure, the 26

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reflections in the area of the Boso SSEs likely result from a thin low-velocity layer along the UPHS, and there is a distinctive convex structure of high velocity structure in the eastern part (X = 85-110 km). A slightly upward high velocity structure may exist at X=45-80 km. By extension of the structure in the adjoining Izu-Bonin arc into the

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region off the Boso Peninsula, these high-velocity structures may represent boninitic

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material of the outer-arc high, partially serpentinized peridotite, or gabbro displaced by

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intraoceanic reverse faults.

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Acknowledgements

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R/V Hakuho-maru of JAMSTEC deployed the OBSs and provided the airgun source

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for the active-source seismic data for this study. The OBSs were retrieved by

Shincho-maru of Shin-Nihon-Kaiji Co. (now Fukada Salvage Co.). We thank the captains and crew of both vessels for their assistance. We also thank Gou Fujie for advice on the application of traveltime mapping and for providing us with the wave-display program Pasteup version 2.1.6. We are grateful to anonymous reviewers for thorough and helpful reviews. Figures in this paper were drawn with Generic Mapping Tools (Wessel and Smith, 1998).

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Fundings: This study was financially supported by the Japanese Ministry of Education, Culture, Sports, Science and Technology under its Observation and Research Program for Prediction of Earthquakes and Volcanic Eruptions, and by a Grant in Aid for

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Scientific Research (25287109).

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All figures are required color in print 2-column fitting image

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Fig. 1 (a) Plate tectonic setting around the Japanese islands. PHS, Philippine Sea plate; PAC, Pacific plate. Orange rectangle is the area shown in (b). (b) Bathymetric map off the Boso Peninsula showing significant tectonic and seismic features. Bathymetric contour interval is

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500 m. Locations of seismic survey lines are shown in red, green, pink, and blue bold lines, which are lines M, S1, S2, and S5, respectively. The line M is the survey line in Kono et al.

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(2017). Yellow triangles show locations of the OBSs. Black triangles are the OBSs which could not retrieved data. Dashed white lines indicate plate boundaries. Closed curves outline source regions of the 1703 Genroku Kanto earthquake (purple line; Sato et al., 2016), 1923 Taisho Kanto earthquake (black line; Matsu’ura et al., 2007), and main slip area of the 2013– 2014 Boso SSE (dashed blue line; Sato et al. 2017). White arrows indicate the directions of motion of the PAC and PHS relative to the landward plate (Seno et al., 1993; Seno and Sakurai, 1996). Black straight line is the survey line of Kimura et al. (2009); black and white bars indicate strong and weak reflectivity area at the upper surface of the PHS, respectively. Pink thin line is the survey line of Nakahigashi et al. (2012). Dashed brown lines are depth contours of the UPHS (Kono et al., 2017). Dashed green line of 20 km contour is from Hirose et al. (2008), western green line of 10 km contour is from Tsumura et al. (2009), and eastern green line of 10 km contour is from Takeda et al. (2007) and Iwabuchi et al. (1990). Orange rectangle is the area shown in Fig. 2.

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2-column fitting image Fig. 2 Map of the study area, showing the coordinate axes of this study’s 3D model (red lines) and survey lines S1, S2, S5, and M. Bathymetric contour interval is 250 m. Dashed brown lines are depth contours of the UPHS (Kono et al., 2017). Dashed green line of 20 km contour is from Hirose et al. (2008), and western green line of 10 km contour is from

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Tsumura et al. (2009). Numbered yellow triangles are OBS stations.

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1.5 column fitting image

Fig. 3 Examples of seismograms recorded by OBSs on the survey line (top) and off the

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survey line (bottom), showing first arrivals (red), later phases (purple with number),

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synthetic first traveltimes (light blue) from the final model and synthetic reflection traveltimes (green) from the UPHS. Vertical axis is the elapsed time from each shot,

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and the reduction velocity is 6.0 km/s. Horizontal axis is the offset from the OBS. A bandpass filter of 5–10 Hz and geometric gain control have been applied. Wave

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2 column fitting image

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amplitudes are normalized by the amplitude of first arrivals.

Fig. 4 2D P-wave velocity structures used for constructing the 3D initial velocity structure. (a) Bathymetric map of the survey area and locations of 2D velocity structures. Line 2 adopted the 2D velocity structure of Kono et al. (2017). (b) Vertical profiles at locations shown on (a). Velocity contour interval is 0.5 km/s. Modeled tectonic features are labeled.

2 column fitting image Fig. 5 Velocity perturbations added to the final velocity model for the checkerboard test. (a) Horizontal sections of added velocity perturbations at 7, 10, 15, and 19 km depth. (b) Vertical section at Y = 35 km, and at X = 85 km.

2 column fitting image Fig. 6 Final 3D seismic velocity structure. (a) Horizontal sections at 5, 9, 13, and 17 km 37

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depth with 8, 10, 15 and 20 km contours of the UPHS. Black dashed lines indicate the UPHS by previous studies (Kono et al., 2017, Hirose et al., 2008, and Tsumura et al., 2009). Blue dashed lines indicate extended UPHS by this study. Blue dot lines show eastern extended area of the UPHS. Brightened areas are parts of the model with ray paths, and darkened areas are those without ray paths. Velocity contour interval is 0.5 km/s. (b) Vertical sections at Y = 35 km and X = 60 km. (c) Vertical sections at Y = 50 km and X = 90 km. Solid lines represent the seafloor. Dashed lines indicate the UPHS.

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Fig. 7 Result of the checkerboard test showing inverted results of the velocity perturbation

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pattern of Fig. 5. (a) Horizontal sections of the inverted result at 7, 10, 15, and 19 km depth. Black dashed lines indicate the UPHS by previous studies (Kono et al., 2017, Hirose et al.,

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2008, and Tsumura et al., 2009). Green dashed lines indicate extended UPHS by this study.

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Green dot lines show eastern extended area of the UPHS. (b) Vertical section at Y = 35 km

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and X = 85 km.

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Fig. 8 (a) Plan view of 3D traveltime mapping results. Red lines are depth contours of the UPHS. White bars with number indicate mapping areas of the corresponding later phases in

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Fig. 3. (b), (c) Oblique views of 3D traveltime mapping results shown in multiple vertical sections. Gray clouds represent the position and uncertainty of reflections. Solid black lines represent the seafloor; dashed red lines represent the UPHS. (b) Cross sections at X = 15, 30, 45, 60, 75, and 90 km. Arrows with number indicate the traveltime mappings of the corresponding later phases in Fig. 3. (c) Cross sections at X = 20, 35, 50, 65, 80, and 95 km.

2 column fitting image Fig. 9 (a) Plan view of 3D travel time mapping results within ±2.5 km depth from the UPHS. Red lines are depth contours of the UPHS. Blue dot lines show eastern extended area of the UPHS. Light blue lines indicate reflection parts within ±2.5 km depth from the UPHS implied by Kono et al. (2017). (b) Observable range of reflection traveltimes from the UPHS, signified by black shading. See the text for details.

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1 column fitting image Fig. 10 (a) Plan view of 3D traveltime mapping results within ±2.5 km depth from the UPHS (same as Fig. 9a) plus results of previous studies. Dashed blue line outlines the main slip area of the 2013–2014 Boso SSE (Sato et al. 2017). Straight diagonal line is the survey line of Kimura et al. (2009); black and white segments represent areas of strong and weak reflections on the UPHS, respectively. Inverted red triangles are the positions of ocean bottom pressure gauges reported by Sato et al. (2017). Dashed orange lines show locations of vertical sections at (b) and (c). (b) and (c) Vertical sections of

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traveltime mapping results combined with the final velocity structures at (b) Y = 35 km, and (c) X = 40 km. Dashed lines represent the UPHS. Dashed light blue ellipses indicate the

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structure. Velocity contour interval is 0.5 km/s.

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reflective area on the UPHS. Blue arrow in (b) indicates a slightly upward high velocity

1. column fitting image

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Fig. 11 (a) Plan view of 3D traveltime mapping results within ±2.5 km depth from the UPHS plus results of previous studies. Areas between dashed brown lines, and dashed blue lines are the projected extension of the outer-arc high in the Izu-Bonin arc for straight

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PHS-PAC boundary case, and deformed boundary case, respectively. Dashed light blue line is the northern limit of the PHS estimated by Uchida et al. (2009). Dashed green outline

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is the location of the high-velocity structure of Ito et al. (2017b) at 15 km depth. Dashed orange line shows locations of vertical section at (b). (b) Vertical section of traveltime mapping result combined with the velocity structure at Y = 39 km. Dashed lines represent the UPHS. Dashed light blue ellipse indicates a convex shape of high velocity with a high velocity gradient in the PHS. Blue arrow indicates a slightly upward high velocity structure. Velocity contour interval is 0.5 km/s.

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