Download Ambient seismic noise tomography of the southern East Sea (Japan

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Ambient seismic noise tomography of the southern East Sea (Japan Sea) and the Korea
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Strait
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: Group velocity maps for the southern East Sea and the Korea Strait with three distinct velocity
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anomalies
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Sang-Jun Lee1,a, Junkee Rhie1,b,*, Seongryong Kim1,2,c, Tae-Seob Kang3,d and Gi Bom Kim1,e
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Korea.
School of Earth and Environmental Sciences, Seoul National University, Seoul 151-742,
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Research School of Earth Science, Australian National University, Canberra, ACT, Australia
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Department of Earth and Environmental Sciences, Pukyong National University, Busan 608-
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737, Korea.
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a
[email protected] , b [email protected], c seongryong,[email protected], d [email protected],
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e
[email protected]
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*
Corresponding author : Junkee Rhie
E-mail: [email protected]
Tel: +82-2-880-6731
School of Earth and Environmental Sciences
Seoul National University
Seoul 151-742, Korea
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Abstract
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Group velocity maps were derived for the southern East Sea (Japan Sea) and the Korea Strait
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(Tsushima Strait) for the 5-36 s period range, which is sensitive to shear wave velocities of the
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crust and the uppermost mantle. Images produced in our study enhance our understanding of
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the tectonic evolution of a continental margin affected by subducting oceanic slabs and a
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colliding continental plate. The seismic structure of the study area has not been described well
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because seismic data for the region are scarce. In this study, we applied the ambient noise
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tomography technique that does not rely on earthquake data. We calculated ambient noise
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cross-correlations recorded at station pairs of dense seismic networks located in the regions
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surrounding the study area, such as the southern Korean Peninsula and southwestern part of
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the Japanese Islands. We then measured the group velocity dispersion curves of the
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fundamental mode Rayleigh waves from cross-correlograms and constructed 2-D group
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velocity maps reflecting group velocity structure from the upper crust to uppermost mantle.
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The results show that three distinct anomalies with different characteristics exist. Anomalies
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are located under the Ulleung Basin (UB), the boundary of the Basin, and the area between
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Tsushima Island and the UB. 1-D velocity models were obtained by inversion of dispersion
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curves that represent vertical variations of shear wave velocity at locations of three different
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anomalies. The 1-D velocity models and 2-D group velocity maps of lateral variations in shear
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wave group velocities show that the high velocity anomaly beneath the UB originates from
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crustal thinning and mantle uplift. Confirming the exact causes of two low velocity anomalies
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observed under the UB boundary and between Tsushima Island and the UB is difficult because
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additional information is unavailable. However, complex fault systems, small basins formed
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by faulting, and deep mantle flow can be possible causes of the existence of low velocity
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anomalies in the region.
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Keywords: East Sea, Ulleung Basin, Korea Strait, Ambient noise tomography, Group velocity
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1. INTRODUCTION AND BACKGROUND
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There have been many previous studies focusing on relatively large-scale features in the upper
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mantle of the northeastern region of Asia (Huang and Zhao, 2006; Zheng et al., 2011; Wei et
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al., 2012). Along with the large-scale studies, detailed research on inland regions has been
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performed using data recorded at dense seismic networks in the Japanese Islands (JI) and
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Korean Peninsula (KP) (Kang and Shin, 2006; Yoshizawa et al., 2010; Guo et al., 2013; Liu et
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al., 2013). The seas between the KP and JI are of particular interest to seismologists because
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surrounding plate interactions in the past triggered multiple tectonic events (Fig. 1., Chough
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and Barg, 1987; Tamaki et al., 1992; Kim et al., 2008; Son et al., 2013). However, the detailed
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seismic structures have not been understood well and only limited information from active-
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source seismic experiments is available. Lee et al. (2001) proposed a seismic stratigraphic
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interpretation of the Ulleung Basin (UB) by analyzing multi-channel seismic reflection profiles.
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Lee et al. (1999) and Kim et al. (2011) revealed the crustal structure and volcanic history of
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the UB. Cho et al. (2004), Sato et al. (2006), and Lee et al. (2011) determined the marginal
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structure of our study region and presented possible tectonic causes. These studies yielded well-
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resolved shallow (sediment-to-crustal) structures along profiles. But most of profiles sampled
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the UB, while relatively few profiles sampled marginal areas, limiting our understanding of the
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region. The maximum resolvable depth of these studies is near the depth of the Moho, further
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limiting the applicability of existing research.
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Passive seismic methods can help overcome these limitations. Due to a lack of seismic stations
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on the ocean floor, methods that use body waves to reveal seismic structures are not considered
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appropriate. Therefore, surface wave method can be used as an alternative to study seismic
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velocity structures in a given region. Recent surface wave studies in the southern East Sea (ES)
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and Korea Strait (KS) help characterize our smaller study area. Yoshizawa et al. (2010) used
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interstation dispersion data measured from earthquake waveforms and Zheng et al. (2011) used
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surface wave dispersion data obtained from ambient noise. Both investigations used data
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recorded at multiple stations in Northeast Asia to generate high-resolution images beneath the
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southern ES and KS. However, the studies utilized only one station in the KP. By using multiple
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stations in the KP, we can enhance the resolution of the models for our study area.
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Relatively long inter-station distances potentially restrict use of high-frequency dispersion data
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for studying the shallow crustal structure of the southern ES and KS. Lack of permanent
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seismic stations within the study area also potentially limit investigation of deeper structures
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in the basin. However, the density of the seismic network surrounding the study area is very
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high, even though azimuthal coverage is not perfect. Ambient noise tomography methods can
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address coverage problems in this situation because station pair paths create an intersecting
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network beneath the study area. Ambient noise tomography has been successfully applied in
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Korea and Japan before (Kang and Shin, 2006; Guo et al., 2013). Inter-station distances in the
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target region are a possible concern because of potential problems extracting Green’s functions.
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Lin et al. (2006) addressed the concern by showing it is possible to extract Green’s functions
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by cross correlating station paths beneath the ocean. Also, feasibility tests verified that
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tomographic methods using array data surrounding oceans are effective (Pawlak et al., 2011;
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Zheng et al., 2011). Therefore, we selected ambient noise tomography to study group velocity
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structures and lateral seismic velocity changes in the region.
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2. TECTONIC SETTING OF THE EAST SEA AND KOREA STRAIT
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East Asia is located at the eastern margin of the Eurasia plate, in a tectonic setting complicated
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by the subducting Pacific and Philippine Sea plates and the colliding India plate. The India–
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Eurasia continental collision and the East Asian subduction have created many back-arc basins
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and Cenozoic normal and strike-slip faults in the study region (Schellart and Lister, 2005; Wei
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et al., 2012). Subducting plates cause complex and various tectonic activities such as
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earthquakes in and around the ES and volcanism along the JI, and numerous marginal basins
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are formed as a result of slab rollback processes (Schellart and Lister. 2005). Beneath the
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Kyushu region, a low velocity anomaly exists that appears to represent upwelling from a mantle
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wedge disturbed by a subducting slab (Seno, 1999; Sadeghi et al., 2000; Yoshizawa et al.,
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2010). The hot upwelling corresponds to volcanic chains in the Kyushu region, including the
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Aso volcano (Zheng et al., 2011).
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The ES is a back-arc basin that started to form in the late Oligocene (Tamaki et al., 1992). Three
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deep basins (Japan, Yamato, and Ulleung basins) and some topographic highs (Korean plateau,
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Yamato rise, and Sea Mountains) are the main features of the ES. The basins were formed
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during tectonic extension, and elevated areas are remnants of continental crust or volcanic
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deposits. After the ES opened, the KS formed by clockwise rotation of southern Japan caused
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by northward motion of the Philippine Sea plate 15–16 Ma (Kim et al., 2008; Son et al., 2013).
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Crustal thickness of the ES and KS varies from place to place. The crust of the ES is generally
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thicker than normal oceanic crust (~7 km) and the thickness at UB is more than double the
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average (~20 km). In contrast, the KS consists of relatively thick continental crust (~30km)
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with a relatively thin sediment layer. Crustal thickness is greatest at the KP and JI (~36 km)
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(Chough and Barg, 1987; Cho et al., 2004; Gil’manova and Prokudinf, 2009; Kim et al., 2011).
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Many fault systems exist around the UB and KS. Along the western margin of the UB, there
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are the Hupo and Ulleung faults and associated small folds and thrust faults. The Hupo fault
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running along the continental slope of the eastern Korean continental margin is a long, narrow
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normal fault (~140 km long and ~3–4 km wide) with vertical displacement up to 1 km. The
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Ulleung fault is a strike-slip fault running N-S to NNE-SSW (Kang et al., 2013; Yoon et al.,
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2014). The Tsushima-Goto fault system is a major tectonic line around the Tsushima Island
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(TI), with normal fault systems dominating the western part, and thrust faults and folds
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dominating the eastern side. The area also contains compressional structures like the Dolgorae
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Thrust Belt (DTB) northeast of TI, and the San`in fold belt between TI and JI (Kim et al., 2008;
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Son et al., 2013).
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3. DATA COLLECTION AND MODELING METHODS
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Continuous waveforms recorded at 60 broadband seismic stations in 2008 were used for our
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study. We obtained vertical component waveforms from 28 stations in the KP deployed by the
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Korea Meteorological Administration (KMA) and the Korea Institute of Geoscience and
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Mineral Resources (KIGAM). In addition, we used continuous waveforms from 31 stations
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belonging to the F-net broadband seismograph network operated by the National Research
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Institute for Earth Science and Disaster Prevention (NIED) and from 1 station of the Global
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Seismic Network (GSN) (Fig. 1).
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To understand group velocity structure in the study area, we constructed group velocity maps
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for several periods (5, 10, 20, 25, 30, and 36 s) using given waveform data. First, we calculated
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the noise cross-correlation function (NCF, which represents Green’s function of each station
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pair) for all possible station pairs. Second, the group velocities of fundamental mode Rayleigh
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waves were measured for different periods with a 1 s increment. Finally, theoretical travel times
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for all station pairs were inverted for 2-D group velocity maps of specific periods. The NCF
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representing seismic wave propagation between two stations can be extracted from the cross-
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correlation of ambient noise recorded at two given stations (Shapiro and Campillo, 2004;
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Shapiro et al., 2005; Bensen et al., 2007). This method is only valid when the ambient noise
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consists of diffusive wave fields. Specifically, the temporal and spatial distribution of the noise
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sources should be random. Therefore, reducing effects of transient but coherent signals (e.g.,
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earthquakes) is important, and two methods are widely used. The first method normalizes
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amplitudes of the waveforms in the time domain. We utilized the second method, which
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removes the portion of waveforms likely contaminated by earthquakes or unknown localized
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sources.
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NCFs were extracted using the following procedures: (1) Continuous vertical component
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waveforms were divided into 1 h windows with 30 min overlap. (2) The mean, trend, and
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instrument response of each window were removed. (3) Sampling rates of waveforms were
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decreased to 1 sample per second (sps) if the sampling rate of the original waveform was not 1
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sps. (4) A band pass filter between 0.01 and 0.45 Hz was applied. (5) The windows that
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appeared to be contaminated by earthquakes were removed. We discarded windows if their
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peak amplitudes were 10 times larger than the median root mean square (RMS) of all windows
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for each station. (6) Cross-correlation was calculated for all possible pairs of 1 h windows. In
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this step, spectral whitening was applied by dividing a 40-point moving average in the
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frequency domain. (7) The final correlogram for each station pair was calculated by averaging
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all remaining correlograms that satisfied the condition in step (5). (8) The NCF was extracted
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from the correlogram. The correlogram consists of causal (positive lag) and acausal (negative
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lag) sections. The two sections represent the NCFs from one station to another. When ambient
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noise is completely diffusive, both sections appear identical. In many cases, however, the
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distribution of noise sources is not completely random. Therefore, we must pick the optimal
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NCF for a given station pair between the causal section, acausal section, or average of two
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sections. Strong and continuous seismic sources are known within the study area, and they
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include the well-documented Kyushu microseisms that are explained by activity of the Aso
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volcano. The detailed procedures for reducing the effects of the Kyushu microseisms are
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described in a previous study (Zheng et al., 2011). We used the same procedures, but took the
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location of the Aso volcano as the source location of the Kyushu microseisms. We manually
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checked the correlograms and only included NCFs that were less contaminated by the Kyushu
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microseism activity.
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To measure group velocities from the NCF, the multiple filter technique of Herrmann and
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Ammon (2002) was applied. We filtered the NCF using a narrow Gaussian filter with a specific
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center period and calculated an envelope. By repeating this process for different periods we
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constructed Frequency Time Analysis (FTAN) diagrams, which present energy distribution as
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a function of time and period (Fig. 2). The group velocity dispersion curve of the fundamental
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mode Rayleigh wave can be obtained from FTAN diagrams by picking corresponding peaks at
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each period. The group velocity was measured only if the inter-station distance of the station
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pair was 3 times longer than the wavelength at the given period. To reduce possible errors, we
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manually measured the group velocities when the corresponding peaks were clearly identified
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in the diagram. The total number of group velocity measurements varied with the periods, but
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more than 1000 velocities were measured in all cases. The average group velocity dispersion
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curve of the region was calculated by bootstrapping, and the average values were used as
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reference values for 2-D group velocity maps.
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A nonlinear 2-D tomographic method was applied for the inversion. The propagation paths
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between the stations were updated by using the fast marching method (FMM) (Sethian and
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Popovici, 1999; Rawlinson and Sambridge, 2005; Rawlinson, 2005). The subspace method
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(Kennett et al., 1988) was used to find optimal model perturbations in each iteration. The
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combination of FMM and the subspace method has been verified to provide stable results in
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many previous studies (Rawlinson and Sambridge, 2005; Saygin and Kennett, 2010; Kim et
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al., 2012). We selected optimal parameters for damping and smoothing by choosing an
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inflection point in the trade-off curve between the RMS misfit and the model variance for each
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period. The final damping parameter varies from 10 to 20 for different periods and the
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smoothing parameter is fixed to be 10 for all periods.
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4. RESULTS AND DISCUSSION
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4.1. Reliability and verification
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The group velocity maps reflecting structure beneath the study region have been developed for
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periods at 5, 10, 20, 25, 30, and 36 s (Fig 3). To evaluate the reliability of the maps, we first
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conducted checkerboard tests. We constructed testing models with three different checker sizes
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(0.6°x0.6°, 1°x1° and 1.3°x1.3°) with ±10% velocity perturbations. Synthetic arrival time data
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were produced for the same distributions of inter-station paths with observations. The synthetic
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data were inverted for the group velocity maps using the same inversion scheme and the
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resulting images were compared with the original testing models (Fig 4). The test results show
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that checkers with even the smallest size are well recovered in the KP and JI where the ray
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density is high and azimuthal coverage is good (Fig 4). In the ES and the KS, structures with
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sizes greater than 1°x1° are resolvable. However, the recovered images using smaller checkers
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are smeared in the NW-SE direction around the ES and the KS. The pattern implies that small-
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scale anomalies shown in the resulting group velocity maps could be artifacts or may be
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extended in the NW-SE direction.
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To verify the reliability of our results, the group velocity maps were compared with the
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previous results for the southern KP and the southwestern JI. A significant low velocity
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anomaly is observed in the Kyushu region of JI in this study, consistent with an ambient noise
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phase velocity map constructed from a previous study (Guo et al., 2013). The low velocity
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anomaly at Kyushu coincides with volcanic regions at the surface. In contrast, group velocity
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structures obtained in this study around the Shikoku and the Nankai regions look quite different
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from the results of Guo et al. (2013). The discrepancy may relate to poor resolution of our map
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in these regions, as shown in our checkerboard tests. For the KP, the absolute level of velocity
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perturbation is less than 4% for all periods, indicating that lateral variation of seismic structures
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beneath the KP are not significant. Previous results of ambient noise tomography in the KP
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highlight shallow crustal structure (Kang and Shin, 2006; Cho et al., 2007, Choi et al., 2009)
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and show that velocities at the northwestern and the southeastern parts of the KP are lower than
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other places. A similar pattern is observed in our group velocity map for the 5 s period (Fig 3a).
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4.2. Regional velocity anomalies
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The group velocity maps of different periods show three distinctive anomalies beneath the
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southern ES and KS (Fig 3). Anomaly 1 is a significant high velocity zone beneath the UB for
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periods between 20 and 30 s. Anomaly 2 is a low velocity region observed at 5 and 10 s periods
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in continental shelves east of the KP and in the northwest of the southwestern JI, surrounding
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the UB. Anomaly 3 is the most prominent low velocity anomaly around the KS, and clearly
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appears at much wider period ranges. The three anomalies are identified based on lateral
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variation of group velocities for different periods. A 1-D shear wave velocity model was
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calculated from our measured average dispersion curve. The model shows Rayleigh wave
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group velocities for 5 and 36 s periods are highly sensitive to shear wave velocities at depths
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of 5 and 40 km, respectively (Fig. 5). Applying this information, the calculated depths of
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Anomalies 1, 2, and 3 are roughly 18–30, 5–8, and 5–40 km, respectively.
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In order to estimate more exact variations of shear wave velocities with depth, we performed
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1-D depth inversions using local dispersion curves of group velocity for 4 different locations.
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We selected locations to characterize and compare regions of velocity anomalies shown in Fig.
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3: the basement of the UB (Fig. 6a) for Anomaly 1, continental shelves around JI (Fig. 6b) for
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Anomaly 2, the center of the low-velocity anomaly in the KS (Fig. 6c) for Anomaly 3, and the
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southern sea of the KP (Fig. 6d) as a reference. The local dispersion curves were created by
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spline interpolation using the group velocity maps in the range of 5 to 36 s periods with 1 s
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increments (Fig. 3). The iterative inversion method of Julia et al. (2000) was utilized to obtain
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shear velocity models. For the starting model of each inversion, we used slightly modified
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models from CRUST1.0 to allow good convergence. After several trials using smoothing
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values in the 0-10 range, we constrained the smoothness to 1.0 for the inversions. Synthetic
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dispersion curves fit the observed local dispersion curves with root-mean-square errors less
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than 0.025 km/s (Fig 6). We produced 4 different 1-D models representing depth variations of
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shear wave velocity at 4 locations and call them Model 1 through Model 4.
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Possible causes of each velocity anomaly are varied. For Anomaly 1, the lateral extent of the
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anomaly is confined within the UB. The related Model 1 has greater velocity than reference
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Model 4 over all depth ranges. Because surface wave dispersion is smoothly sensitive to shear
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velocities across depths (Fig. 5), the increased velocities and rapid velocity changes with depth
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likely indicate a shallower crust/mantle transition. Previous observations of thinner crust in the
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region support this explanation and the thinner crust is explained by extension following the
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opening of the ES (Kim et al., 1998; Sato et al., 2006; Park et al 2009; Lee et al., 2011). We
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infer from our model and supporting study findings (Zheng et al., 2011) that Anomaly 1 results
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from locally uplifted mantle structure. At the location of Anomaly 1, the effect of the water
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layer may not be negligible because the depth of water is around 2 km (e.g., Huang et al., 2014).
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To check the reliability of the features resolved by the inverted model for this location, we
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performed an inversion with a water layer as well (Herrmann and Ammon, 2002). For this
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purpose, the previously inverted model was taken as the initial model. Since modeling the
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dispersion curve at short periods (5-7 s) gives numerically unstable results, data for periods
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longer than 7 s were used. The results showed that the inverted model with the water layer was
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comparable with the previous model, except for the shallow depth (~15 km) (Fig 6a). However,
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the increased velocity in the upper part of the newly inverted model provided additional
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evidence in support of the described features, in spite of the introduced water layer modifying
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the model. For other anomalies, the effect of the water layer was negligible because the
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thickness of the water layer was less than a couple of hundred meters.
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For Anomaly 2, the related Model 2 shows relatively lower velocity than reference Model 4 at
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depths shallower than 10 km. In map view, Anomaly 2 surrounds the UB and appears only in
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the ocean. Anomaly 2 may represent distribution of sediments, but the sediment layers along
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the eastern Korean continental margin are thinner than sediments within the UB (Chough and
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Barg, 1987; Sato et al., 2006; Gil’manova and Prokudin, 2009; Park et al 2009; Lee et al., 2011).
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Other possible explanations for low velocity anomalies at shallow crust of the UB boundary
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are faults and related structures, such as the Ulleung, Hupo, and Yangsan faults along the
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eastern coast of the southern KP and western margin of the UB. Faults and folds are also
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distributed between Tsushima and Oki islands (Kim et al., 2008; Gil’manova and Prokudin,
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2009; Yoon et al., 2014).
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In map view, Anomaly 3 appears inside Anomaly 2 and the velocity of Model 3 is slower than
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that of reference Model 4 in the crust and upper mantle. Anomaly 3 and Anomaly 2 may share
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the same origin at shallow depths, but Anomaly 3 could extend deeper into the upper mantle.
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Therefore, similar to Anomaly 2, sedimentary structures are not the likely cause of the anomaly.
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Significant fold and fault systems near Anomaly 3, such as the Tsushima-Goto fault, Dolgorae
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Thrust Belt, and San’in folded zone (Kim et al., 2008; Son et al., 2013; Yoon et al., 2014),
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could be an origin of the low velocity anomaly in the shallow crust. However, the fold and fault
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systems do not likely explain the low velocity anomaly extending to the upper mantle. A
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complementary component that may explain the deep low velocity anomaly is mantle
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upwelling flow. A previous study reported that low velocity anomalies exist in the upper mantle
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beneath the eastern and western boundaries of the ES and they indicate upwelling flow, from
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slab dehydration at several depths, possibly caused by phase transitions and stagnation of the
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slab (Zhao et al., 2007; Zheng et al., 2011). The surface trace of the upwelling flow reported
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by Zheng et al. (2011) closely matches the low velocity region within our study area. This
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mantle upwelling may facilitate hydrothermal fracturing. From this perspective, Anomaly 3
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can be explained by significant fold and fault systems that are the local center of enhanced
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hydrothermal activities originating from upwelling mantle flow. The lack of an obvious low
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velocity anomaly in the upper mantle corresponding to Anomaly 2 may be because of low
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resolution of long period surface waves that sample the upper mantle, or because of relatively
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weak hydrothermal activities in the given area.
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The relationship between the group velocity structure and seismicity supports our low velocity
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anomaly hypothesis. We related the group velocity map at 5 s and seismicity from the ISC-
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GEM Catalog (Storchak et al., 2013) (Fig. 7). Compared to the UB and the surround marginal
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area, seismicity is clearly more focused in the low velocity region. Generally, shear wave
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velocity is negatively correlated to the thermal structure and amount of melt in the crust. In
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contrast, seismicity is lower in thermally enhanced regions (Kim et al., 2012). However, high
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seismicity occurs in the low velocity region and low seismicity occurs in the high velocity
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region of our study area, implying that the seismicity pattern in our study area cannot be
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explained by thermal structure only. Even if the pre-existing fault systems are more developed
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in the KS region (Son et al., 2013; Yoon et al., 2014), they simply do not explain the high
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seismicity in the predominately low velocity region. A suggested cause of weak crustal material
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and increased seismicity could be water that is continuously supplied to the crust by mantle
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upwelling flow (Hasegawa et al., 2005). We speculate that water supplied by upper mantle
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upwelling enhances seismic activity while reducing seismic velocity by increasing melts and
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heat.
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5. CONCLUSION
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To study seismic group velocity structure of the southern ES and KS, 2-D group velocity maps
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and 1-D shear wave velocity models were calculated by using ambient noise data. Our
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observations show that a significant high velocity anomaly exists in the UB and low velocity
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anomalies of different periods exist in marginal areas. The high velocity anomaly can be
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explained by locally uplifted mantle structure beneath the UB. We can conclude that the crustal
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thickness of the UB is relatively thin compared to surrounding regions. In contrast, the shallow
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crust of the UB boundary has a low Rayleigh wave group velocity and a localized low velocity
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anomaly extends from the crust to uppermost mantle in the region between Tsushima and the
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UB. We suspect that the low velocity anomalies described in our study might be associated
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with hydrothermal activities of mature fold and fault systems in the area. Comparing our group
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velocity map with regional seismicity, we demonstrate that zones with high seismicity usually
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have low velocities. This observation may be explained by weakened crustal strength combined
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with a reduced seismic velocity caused by water supplied from the upwelling mantle.
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Acknowledgement
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We thank the operators of KMA, KIGAM and NIED for making their data publicly available.
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This work was funded by the Korean Meteorological Administration and Development
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Program under Grant CATER 2012-5051.
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Figures captions.
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Fig 1. (a) Topographic map of the East Sea, South Sea and surrounding regions. Some intra-
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volcanoes (yellow triangles) and many subduction zone volcanoes (red triangles) exist in this
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area. The yellow contour indicates subduction depth of the Pacific plate and Philippine Sea
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plate. Features: KP, Korean Peninsula; JI, Japan Island; ES, East Sea (Sea of Japan), KS, Korea
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Strait; YS, Yellow Sea; NT, Nankai Trough; JT, Japan Trench; UB, Ulleung Basin; JB, Japan
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Basin; YB, Yamato Basin; TI, Tsushima Island. (b) Location of KMA, KIGAM, F-net and GSN
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stations (green triangles) and all possible ray paths (gray lines) used in this study.
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Fig 2. Average group velocity for each period (white dot) is obtained from the average
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Frequency Time Analysis (FTAN) diagram using the FTAN method. Amplitude of this figure
517
represents average amplitude of each FTAN diagram, normalized by the maximum amplitude.
518
The group velocity dispersion curve of the fundamental mode Rayleigh wave obtained from
519
the FTAN diagram is used as reference velocity for the 2-D group velocity map and to calculate
520
the sensitivity kernel for depth.
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Fig 3. Group velocity maps (a–f) represent velocity perturbations from the average group
523
velocity for each period at 5, 10, 20, 25, 30, and 36 s. Significant high velocity anomalies are
524
clearly seen beneath the UB for periods between 20 and 30 s (Anomaly 1). Low velocity
525
anomalies surrounding the southern part of the basin are observed at 5 and 10 s (Anomaly 2).
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Localized low velocity anomalies between TI and the UB appear at all periods (Anomaly 3).
527
Additionally, 1-D depth inversion is conducted at the white dots (1–4) that represent each
528
anomaly and a reference average.
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24
530
Fig 4. Checkerboard test results with three different checker sizes (0.6°x0.6°, 1°x1° and
531
1.3°x1.3°). Results show that checkers with 0.6°x0.6° are well recovered in the KP and JI.
532
Recovered checker images of all checker sizes seem to be elongated in NW-SE direction
533
beneath the southern ES and KS.
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Fig 5. Sensitivity kernel at 5, 10, 20, 25, 30 and 36s and 1-D structure inverted from the average
536
FTAN diagram. Rayleigh wave group velocities for periods of 5 and 36 s are sensitive to shear
537
wave velocities at depths of 5 and 40 km, respectively.
538
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Fig 6. 1-D depth inversion results for points (1–4) that represent characteristic velocity
540
anomalies and the reference average. The black line represents the initial model of each point
541
and the inverted model is shown as a red line. The green line represents the inversion result of
542
model 4. The blue line indicates the inverted model with a 2-km thick water layer.
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Fig 7. Group velocity map for the 5 s period and earthquake (2 ≤ M ≤ 6) epicenter locations
545
from 1994 to 2013 (ISC-GEM Catalog; Storchak et al., Seismological Research Letters, 2013).
546
Clearly, seismicity inside the UB is remarkably lower than surrounding regions.
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25