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Journal of Geophysical Research: Solid Earth
RESEARCH ARTICLE
10.1002/2014JB011522
Key Points:
• We investigated the structure below
Mount Fuji from a receiver
function analysis
• A gap of the velocity boundary below
Mount Fuji is a weaker velocity contrast
• A boundary at 25 km deep represents
the bottom boundary of the
magma chamber
Correspondence to:
S. M. Kinoshita,
[email protected]
Citation:
Kinoshita, S. M., T. Igarashi, Y. Aoki, and
M. Takeo (2015), Imaging crust and
upper mantle beneath Mount Fuji,
Japan, by receiver functions, J. Geophys.
Res. Solid Earth, 120, 3240–3254,
doi:10.1002/2014JB011522.
Received 10 AUG 2014
Accepted 29 MAR 2015
Accepted article online 3 APR 2015
Published online 7 MAY 2015
Imaging crust and upper mantle beneath Mount
Fuji, Japan, by receiver functions
S. M. Kinoshita1, T. Igarashi1, Y. Aoki1, and M. Takeo1
1
Earthquake Research Institute, University of Tokyo, Tokyo, Japan
Abstract Mount Fuji has ejected a huge amount of basaltic products during the last 100,000 years. Even
though the region around Mount Fuji is tectonically active, the seismicity below Mount Fuji is low, resulting in
little knowledge about the seismic structure there. To gain more insight into the magma-plumbing system,
we obtain the seismic structure beneath Mount Fuji by the receiver function (RF) technique. RFs at seismic
stations around Mount Fuji show positive phases at ~3 and ~6 s, representing the conversion of P to S waves
at a positive velocity boundary in the Philippine Sea plate. Cross sections of RF amplitudes reveal two distinct
velocity boundaries around Mount Fuji, at depths of 40–50 km and 20–30 km, which we interpret to be the
boundary between the crust-mantle transition layer and the uppermost mantle of the Izu-Bonin arc and the
velocity discontinuity just below the region where low-frequency earthquakes (LFEs) of Mount Fuji have
occurred, respectively. The velocity boundary at about 50 km depth shows a clear gap just beneath Mount
Fuji. We suggest that this gap represents a weaker velocity contrast zone through which the magma of Mount
Fuji ascends from the Pacific plate. A thorough grid search reveals that a low-velocity zone at depths of
~13–26 km explains all the characteristics of RFs around Mount Fuji, leading us to interpret the high-velocity
boundary just below the LFE region as the lower boundary of Mount Fuji’s magma chamber.
1. Introduction
Mount Fuji, an arc volcano in central Japan, has at least two unique features. First, the eruption rate of Mount
Fuji has been much larger than that of other island-arc volcanoes in Japan. Tsukui et al. [1986] estimated that
the average eruption rate of Mount Fuji is about 5 km3/ka, while typical subduction-related volcanoes in
Japan are about 0.01–0.1 km3/ka. Second, Mount Fuji has ejected basaltic products during the last
100,000 years, while a typical eruption product from an island-arc volcano is andesitic [Fujii, 2007]. Although
these characteristics suggest that the magma-plumbing system of Mount Fuji is different from typical arc
volcanoes in Japan, an unambiguous image of the magma system has not been produced thus far.
The unique features of Mount Fuji may be due to the complicated tectonics around it. The Philippine Sea
(PHS) plate is subducting beneath the Eurasian plate and the Okhotsk plate along the Suruga and Sagami
troughs, respectively, and the Izu-Bonin arc (IBA), which is on the PHS plate, has been colliding with central
Japan during the last 15 Ma (Figure 1) [Takahashi and Saito, 1997]. Mount Fuji lies on the boundary
between the colliding and subducting regions. Moreover, the Pacific (PAC) plate is subducting from east to
west beneath the PHS plate.
Fujii [2007] noted that basaltic magmas of Mount Fuji have large variations in concentration of incompatible
elements while maintaining constant silica content, which is quite different from other volcanoes along the IBA.
With increasing pressure of magmatic differentiation, the role of pyroxenes increases and the silica increment
with magmatic differentiation is suppressed. Fujii [2007] concluded that the magma reservoir beneath Mount
Fuji is deeper than those beneath other volcanoes along the IBA, lying at a depth of about 20 km or deeper.
Kaneko et al. [2010] analyzed densely piled scoria layers exposed on the middle of the eastern flank of Mount
Fuji. Based on the chemistry of whole rocks, phenocryst, and olivine-hosted melt inclusions, they concluded
that the magma erupted from Fuji is generated through the mixing between basaltic and more SiO2-rich end
members. From this, Kaneko et al. [2010] proposed that Mount Fuji’s magmatic plumbing system consists of
at least two magma chambers, that is, a relatively deep basaltic one at 20 km and a relatively shallow one with
more SiO2-rich end members at 10 km deep.
©2015. American Geophysical Union. All
Rights Reserved.
KINOSHITA ET AL.
Nakamura et al. [2008] carried out geochemical surveys of magmatic rocks from 28 quaternary volcanoes in
central Japan including Mount Fuji. They found that the PHS plate does not seem to significantly inhibit fluid
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Figure 1. The tectonic background around Mount Fuji. (left) Active plate boundary around the Japanese islands. Black lines
represent plate boundaries from Seno et al. [1996]. (right) Tectonic setting around Mount Fuji. Depth contours indicate the
upper surface of the subducting Philippine Sea (PHS) plate estimated by Nakajima et al. [2009]. Black crosses represent
hypocenters with a magnitude 0.1 or larger at depths more than 20 km, between 2001 and 2005 located by JMA. Gray
arrows show the relative plate motion between the PHS plate and the Eurasian plate [Seno et al., 1993].
flow from the PAC plate below in regions where PAC and PHS plates overlap, and the contribution of PHS fluid
is distinctly small in Mount Fuji compared with other areas.
To understand the magma-plumbing system of Mount Fuji, we have to clarify the seismic structure there.
Various studies have explored the structure beneath Mount Fuji from hypocenter distributions, focal
mechanisms, three-dimensional seismic velocity structures, and reflection/refraction surveys. Nakajima et al.
[2009] determined the distribution of interplate earthquakes relocated by an appropriate one-dimensional
velocity model and carried out traveltime tomography to estimate the three-dimensional seismic velocity
structures in central Japan. Their results suggest that the PHS plate subducts to a depth of 140 km to the
northwest of the Izu collision zone. However, they did not find continuous high-velocity anomalies
representing the PHS slab just below Mount Fuji.
A north dipping interface representing the top of the subducted part of the PHS at 25–35 km depth in the
western parts of the Izu collision zone was revealed based on seismic refraction/wide-angle reflection analyses
from active sources [Arai, 2011; Arai et al., 2014]. These studies concluded that the middle part of the IBA had
been accreted beneath the Honshu arc in the process of the collision and subduction, and the top of the
subducted IBA was the lower crust. Sato et al. [2012] compiled the results of refraction and reflection analyses
from active source data around the Izu collision zone and mapped the upper boundary of the subducting
PHS plate, whose upper and middle crusts have been accreted to the continental crust. Their results show
that the upper boundary of the subducting PHS plate is about 20 km deep, 20 km to the northeast of Mount Fuji.
From September to December in 2000 and from April to May in 2001, many low-frequency earthquakes (LFEs)
occurred at depths of 11–16 km beneath Mount Fuji [Nakamichi et al., 2004; Ukawa, 2005]. In response to them,
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seismic and electromagnetic campaigns were conducted between 2002 and 2005 to find the nature of the
deep low-frequency earthquakes and their structural controls. Aizawa et al. [2004] obtained spatially
continuous, high-resistivity regions from the southwest to the eastern side and from the eastern to the
northeastern side beneath Mount Fuji. They also found a low-resistivity region at the depth of 20–50 km
beneath these high-resistivity regions. They interpreted these high- and low-resistivity regions to represent
the parts of the PHS plate and the magma chamber of Mount Fuji, respectively, concluding that the PHS
plate splits at the eastern side below Mount Fuji. On the other hand, Nakamichi et al. [2007] pointed out the
existence of a magma chamber of Mount Fuji at the depth of 15–25 km based on tomographic imaging of
seismic velocity using a dense seismic array. Additionally, their seismic tomography did not find any tearing
of the PHS plate, contrary to Aizawa et al.’s [2004] study.
Local seismic tomography by Nakamichi et al. [2007] resolves only a shallow portion and does not reveal an
entire image of the magma chamber and the magma-plumbing system underneath it. Nakajima et al. [2009]
did not detect velocity discontinuities just below Mount Fuji with seismic tomography, but the existence of
velocity discontinuities should be examined in more detail in order to reveal the existence of the PHS slab. In
this study, we investigate the structure of the PHS plate below Mount Fuji using a receiver function (RF)
technique. By employing the RF analysis, we take advantages of the sensitivity to velocity discontinuities,
as seen on the subducting plate, over conventional seismic tomography, which inherently has a limited
resolution in imaging velocity boundaries. Our goal is to gain more insight into the volcanic system of
Mount Fuji and to address questions such as why Mount Fuji is so large for an arc volcano.
2. Data and Analysis
A coda wave following a teleseismic body wave arrival consists of mode-converted phases. RF is a time series
estimated by deconvoluting vertical components of teleseismic body wave from the horizontal ones. This
deconvolution process removes the source time functions and the effect of instruments from the
waveform, thereby isolating converted waves [e.g., Ammon, 1991].
RFs are easy to define but difficult to compute in a reliable manner because a raw spectral division is unstable
near spectral holes. Various methods have been proposed to calculate stable and reliable RFs numerically. To
identify Ps converted phases from velocity boundaries below Mount Fuji using high-frequency data, we
estimate RFs with the multiple-taper correlation (MTC) method by Park and Levin [2000]. In the MTC
method, an orthonormal sequence of tapers is designed to minimize spectral leakage. The set of tapers
and the associated eigenspectra can be combined to reduce the variance of the overall spectrum
estimates [Park et al., 1987].
We use the teleseismic waveform data recorded at 159 seismic stations around Mount Fuji operated by five
institutes: Earthquake Research Institute, The University of Tokyo (ERI), Japan Meteorological Agency (JMA),
National Research Institute for Earth Science and Disaster Prevention (NIED), Hot Spring Research Institute of
Kanagawa Prefecture, and Nagoya University (Figure 2). Each station is equipped with a three-component
short-period seismometer or a broadband seismometer. Some stations around Mount Fuji operated only
between October 2002 and June 2005. The orientations of seismometers operated by NIED installed at the
bottom of boreholes are corrected using the result due to Shiomi [2013].
We use a global earthquake catalogue from the United States Geological Survey National Earthquake
Information Center to extract teleseismic events with magnitudes larger than 6.0 and epicentral distance
between 20 and 90° from Mount Fuji for the RF analysis.
Data from October 2002 to June 2005 have been used for RFs. Expected P wave arrival times are first
calculated using the AK135 velocity model [Kennett et al., 1995] and then manually picked. Event selection
is based on the criterion that the number of stations with clear P wave arrivals is equal to or greater than 5
for each event. The number of events at each station is from 15 to 221, depending on the signal-to-noise
ratio and deployment time. Figure 3 shows the distribution of teleseismic events used in this study.
We analyze 80 s waveforms for both pre-event noise and signal. The noise window ends 5 s before the P wave
arrival time, and the signal window begins after the noise window. Preprocessing before calculating RFs
includes mean removal, de-trending, and rotation to a great circle path. A total of 15,501 RFs in both radial
and transverse components are obtained in this analysis.
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3. Examples of
Receiver Functions
Figure 2. Distribution of the seismic stations used in this study. Crosses
represent seismic sites operated by several institutes: ERI, JMA, NIED,
the Hot Spring Research Institute of Kanagawa Prefecture, and Nagoya
University. The triangle represents the summit of Mount Fuji. Red and blue
solid lines show the locations of the cross sections of the RF amplitudes in
Figures 5 and 6, respectively. Four stations indicated by stars are shown in
Figure 4 as examples.
We calculate the composite RFs from a
weighted sum of single-event RFs to
reduce noise. We set bins with a width
of 10° and a spacing of 5° because RFs
vary with the back azimuth with dipping
velocity boundaries [Shiomi et al., 2004].
The weight of each RF is calculated using
inverse variance in frequency domain
[Park and Levin, 2000]. After calculating
composite RFs, we apply a cosine-squared
low-pass filter with a cutoff frequency of
1.0 Hz. We then obtained a total of 5997
stacked RFs. Figure 4 shows an example
of stacked RFs lined up by back azimuth
estimated at four stations, H.MSNH,
H.TOIH, MMS and FUJ, shown by stars in
Figure 2. Positive and negative amplitudes
of the RFs are usually interpreted to be
generated at boundaries with upward
and downward decreasing velocities
(the high- and low-velocity boundaries),
respectively.
Figure 3. Epicenter distribution of the 221 teleseismic events used in this study. Solid lines indicate the epicentral distance
from Mount Fuji.
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Figure 4. Examples of RFs with a cutoff frequency of 1.0 Hz as a function of back azimuth. Locations of these stations are
indicated by stars in Figure 2. These stations are operated by NIED (stations H.MSNH and H.TOIH) and ERI (stations MMS and
FUJ). (left) Radial and (right) transverse RFs, respectively. Vertical and horizontal axes denote the back azimuth and the lag time,
respectively. Note that the radial and transverse plots have the same scale. White arrows denote phases discussed in section 3.
Traveltime tomography [Nakajima et al., 2009] shows that the PHS slab subducts below station H.MSNH from
south to north at a depth of about 30 km (Figure 2) where our RF analysis revealed distinct positive phases at
about 5 s in the radial component, representing a Ps converted phase from the Moho boundary of the
subducting slab (Figure 4a). The arrival times of these Moho phases also vary with back azimuth, with
earlier arrivals by waves traveling from the south (back azimuth of 150–210°) than those traveling from the
north, indicating an existence of a slab dipping from south to north below H.MSNH. At station H.TOIH,
positive phases are seen at 2 and 5 s in the radial component, representing velocity discontinuities in the
continental crust such as the Conrad and Moho boundaries (Figure 4b). At stations to the northeast of
Mount Fuji, such as MMS, the direct P phase shifts apparently toward a positive lag time from zero, with a
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broadened shape, indicating an existence of thick volcanic sediment layers below Mount Fuji [Cassidy, 1992].
Distinct positive phases at both 3 and 6 s in the radial component of MMS suggest sharp velocity increases
below Mount Fuji. On the other hand, RFs of FUJ, southwest of Mount Fuji, do not show any positive phases
at 3 s in the radial component (Figure 4d). Such characteristics may represent differences of the basement
structure between the northeastern and southwestern parts of Mount Fuji, as shown by an active source
seismic experiment [Oikawa et al., 2004]. Examining the RFs more closely, the transverse Ps pulses at 3 and
6 s of station MMS change polarity at 210 and between 255 and 275°, respectively (Figure 4c). Also, the
polarity of transverse pulses of FUJ at 6 s changes at 180, between 55–125 and 255–275°. Such polarity
reversal in transverse components is common if dipping boundaries or anisotropic materials exist [Savage
et al., 2007]. The polarity reversals in Figures 4c and 4d are not explained by either a single two-lobed or a
four-lobed pattern, representing the effect of dipping or anisotropic structures, respectively. This implies that
the polarity reversal that we observed is due to both dipping and anisotropic structures.
4. Cross Sections of the Radial RF’s Amplitude
Many previous studies investigate the depth distribution of velocity discontinuities from RF images using dense
seismic arrays [Shiomi et al., 2004; Tonegawa et al., 2006]. In this study, the ray parameter is calculated for each
event-station pair using the AK135 velocity model [Kennett et al., 1995] with epicentral distance and focal depth.
After correcting for station elevation, the RFs are mapped along its raypath traced in three dimensions through
a flat-layer velocity model. Time-to-depth conversion is achieved with the velocity structure derived by
traveltime tomography around Mount Fuji [Nakamichi et al., 2007]. We set the bin widths to 4 km in the
horizontal direction and 1 km in depth. Each bin was 20 km wide perpendicular to the cross section. If
multiple rays intersect the same bin, we take the average of the RF amplitude over the rays.
Figure 5 shows examples of RF cross sections along the lines in Figures 2a–2g. Before stacking RF amplitudes,
each RF with a cutoff frequency of 1.0 Hz is high-pass filtered using a zero-phase second-order Butterworth
filter with a corner frequency of 0.1 Hz. As shown on the southeast-northwest (SE-NW) cross section in
Figures 5a and 5b, the continuous distribution of positive amplitude indicated by black solid lines suggests
that there is a velocity boundary extending from the southeast to the northwest. Comparing this with the
location of the upper boundary of subducting PHS plate inferred from seismic tomography [Nakajima
et al., 2009] indicated by white dashed lines and the distribution of hypocenter indicated by gray dots, we
conclude that the velocity boundary we found in Figures 5a and 5b represents the Moho boundary of the
subducting PHS slab. Similarly, distinct positive phases between 90 and 120 km (at depths 60–80 km) in
Figure 5g imply the Moho boundary at about 10 km beneath the top of the subducting PHS plate.
Figures 5c–5g show distinct positive phases at a depth of 40 km below the region where the IBA has been
colliding with central Japan. Figure 5d shows that this positive phase does not continue to the area below
Mount Fuji (white arrow in Figure 5d) but another positive phase is seen at 20–30 km, just beneath the
region where LFEs of Mount Fuji have occurred.
To construct more detailed images around Mount Fuji, we change the bin width to 10 km perpendicular to
the direction of the cross section. Figure 6 shows the RF cross sections around Mount Fuji along the lines
A–D in Figure 2 with a cutoff frequency of 1 Hz. RF images have similar features in all directions, with
distinct and broad positive and negative amplitudes at the depths of 0–10 km and 10–20 km, that are
caused by the volcanic sedimentary layers near Mount Fuji. Figure 6 also reveals a distinct gap in the
positive discontinuity at the depth of 40 km as shown by white arrows and a high-velocity boundary at the
depth of 20–30 km below Mount Fuji in all directions, as shown in Figure 5d.
The depth of boundaries found in Figures 5 and 6 have errors caused by the assumption of a flat-layered
structure in depth conversion. As discussed in section 3, the polarity pattern in the transverse component
of RF indicates the existence of dipping and/or anisotropic media below Mount Fuji. Dipping layers below
Mount Fuji cause errors in depth estimation from radial RFs because the arrival times of Ps conversion at
velocity boundaries change by different back azimuths. Shiomi et al. [2004] evaluated the error of the
depth conversion to be about 3 km from the boundary at 50 km depth, with a dipping angle of about 10°.
Anisotropic layers also change the arrival times of Ps conversion in radial RFs.
A forward modeling calculation of RFs, including layers with velocity anisotropy between 3 and 8% in the
mantle wedge and oceanic crust to reconstruct RFs in the subduction zone, shows that the arrival times of
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Figure 5
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Figure 6. Examples of cross sections of the RF amplitude with a cutoff frequency of 1.0 Hz around Mount Fuji along four
blue lines shown in Figure 2. Parameters are the same as those for Figure 5, but the bin width is 10 km perpendicular to
the cross section, with a center of cross sections at Mount Fuji. Regions 3 and 4 indicated by black dashed lines are shown in
Figure 7 as examples. Region 5 indicated by a black dashed line is shown in Figure 10a. White arrows indicate the gap in the
velocity boundary discussed in section 5.4.
Ps converted phases in radial components are mostly constant regardless of back azimuths [Wirth and Long,
2012]. The difference of arrival times between the back azimuths is within 0.5 s, so the error of depth
conversion caused by the anisotropic layer is within 5 km in this study.
5. Discussion
5.1. Evaluation of the Multiples
RF results are contaminated by multiple Ps converted phases (PpPs, PsPs, etc.) at shallow velocity
discontinuities. Arrival times of these multiple phases can be calculated assuming a flat-layered velocity
structure. We can separate Ps and PpPs arrivals because the arrival time of the Ps phase increases with
increasing ray parameter, whereas that for PpPs decreases with increasing ray parameter [Salmon et al.,
2011]. Figure 7 shows an example of RFs stacked according to ray parameters for all the rays that intersect
regions 1 and 2 in Figure 5d and regions 3 and 4 in Figure 6a. Bin widths are 0.006 s/km with a spacing of
0.003 s/km. Predicted arrival times of Ps phases (solid gray lines in Figure 7) are calculated from the
velocity structure obtained by a traveltime tomography by Nakamichi et al. [2007]. In Figure 7a, the arrival
times of focusing positive phases found at the depth of 40–50 km in Figure 5d correspond to that of the Ps
phase from 45 km depth, increasing with increasing ray parameter. Similarly, stacked RFs in Figures 7b–7d
show that delay times of focusing phases increase with increasing ray parameters, supporting the idea
that these phases are not multiples of shallow velocity boundaries.
Figure 5. Examples of cross sections of the RF amplitude with a cutoff frequency of 1.0 Hz along seven red lines in Figure 2.
The color scale of the amplitude is shown at the bottom. Gray dots and white circles in each panel are the hypocenters
with magnitude 0.1 or larger occurring between 2002 and 2005, taken from the JMA catalogue. White circles represent
low-frequency earthquakes (LFEs) near Mount Fuji. Pink circles show relocated hypocenters of LFE below Mount Fuji
[Nakamichi et al., 2007]. The green triangle represents the summit of Mount Fuji projected onto line d. White dashed lines
indicate the upper surface of the PHS plate [Nakajima et al., 2009]. Black solid lines are the estimated velocity boundaries
from this study. Regions 1 and 2 indicated by black dashed lines are shown in Figure 7 as examples. The white arrow
indicates the gap in the velocity boundary discussed in section 9.
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Figure 7. Stacked RFs according to ray parameters whose raypaths intersect (a) region 1 and (b) region 2 in Figure 5, and
(c) region 3 and (d) region 4 in Figure 6. Gray solid lines represent the arrival time of Ps converted waves at the depths of
45 km, 50 km, 50 km, and 26 km, shown in Figures 7a–7d, respectively, derived from the result of Nakamichi et al. [2007].
5.2. Possible Velocity Models Below Mount Fuji
We examine possible velocity models below stations MMS and FUJ in Figure 2 by a grid search approach
(Figure 8). We use a similar method to that employed by Igarashi et al. [2011]. Synthetic seismograms from a
given velocity model are calculated using the propagator matrix method [Kennett and Kerry, 1979], assuming
velocity models with flat isotropic layers. After calculating synthetic RFs by the same method as that used to
calculate the observed RFs, we apply a cosine-squared filter with a cutoff frequency of 1.0 Hz. The horizontal
slowness is set to 0.064 s/km, and the density is estimated using Ludwig’s law [Ludwig et al., 1970]. We use
stacked RFs with the back azimuth range between 120 and 180°, the direction of the subduction of the
Izu-Bonin arc. Observed RFs at MMS and FUJ (black lines in Figures 9a and 9b) have five characteristics: (1) a
wide pulse of the direct P phase (phase A), (2) an apparent shift of the direct P phase from zero time lag, (3) a
large negative amplitude after the direct P phase (phase B), (4) a positive amplitude at about 4 s at MMS and
at 6 s at FUJ (phase C), and (5) a negative amplitude at about 7–8 s (phase D). It is essential to figure out
whether these characteristics are caused solely by high-velocity boundaries or by some low-velocity layers. To
address this question, we examine two kinds of velocity models with (Model 1) and without (Model 2) a
low-velocity layer in the middle crust of the PHS plate (Figure 8a). We set five layers in Model 1 and four
layers in Model 2. Eight model parameters v1, v2, v3, v4, Da, Db, Dc, and Dd in Model 1 and six parameters v1,
v2, v3, Da, Db, and Dc in Model 2 are assumed. Parameters Da Dd represent the depth (km) of the bottom
boundary of each layer. Parameters v1 and v2 v4 represent the S wave velocity (km/s) at the top of Layer 1
and in Layers 2–4, respectively. In Layer 1, the velocity increases from v1 (km/s) to v2 (km/s) by (v2 v1)/Da
steps. Based on the results of traveltime tomography by Nakamichi et al. [2007], we set the Vp/Vs ratio to 2.2
in Layer 1 and the low-velocity layer and 1.73 in other layers. Possible parameter ranges are listed in Table 1
and shown in Figure 8b. We set the following constraints on the possible parameter ranges: the Vs of each
layer can take discrete values with a spacing of 1.0 (km/s) in Model 1 and 0.5 (km/s) in Model 2. The spacing
of possible Da, Db, and Dc is 1.0 km in both models and that of Dd is 2.0 km in Model 1. We also constrain
the average P wave velocity above 60 km depth to be less than 7.0 (km/s) to avoid that all boundaries go to
shallow depths and all of the amplitude is created only by the multiple phases of shallow boundaries.
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Figure 8. (a) Velocity models with (Model 1) and without (Model 2) the low-velocity layer searched in the grid search
modeling of RFs. Parameters Da Dd represent the depth (km) of the bottom boundary of each layer, and v2 v4 represent
the S velocity (km/s) in each layer. The velocity increases from v1 (km/s) to v2 (km/s) by (v2 v1)/Da steps in Layer 1. Possible
parameter ranges are listed in Table 1. (b) Possible velocity models in the grid search. The numbers of velocity models are
2,101,948 for Model 1 and 989,472 for Model 2.
Then, a total of 2,101,948 synthetic RFs for Model 1 and 989,472 synthetic RFs for Model 2 have been
calculated. The amplitudes of all the RFs have been normalized by the amplitude of the direct P phase.
The least squares errors between the synthetic and observed RFs are calculated. Figure 9 shows the
optimum S velocity models for Models 1 and 2 at the two stations. We visualize our result using S velocity,
because the converted wave of RFs is sensitive to S velocity contrast [Julia, 2007].
The optimum solutions we obtained (Figure 9) have the following features:
1. A broadened pulse and apparent shifts from zero time lag (phase A) are explained by setting 1–3 km thick
volcanic sediment layers at both stations. The time shift is larger in the RF of FUJ than that of MMS,
representing a thicker sediment layer on the western part of Mount Fuji than on the eastern part. This
is consistent with the shallow P wave velocity structure derived from an active source seismic experiment
[Oikawa et al., 2004].
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Figure 9. (a) Comparison between synthetic and observed RFs. The solid line represents observed RFs at station MMS in
Figure 2. Gray solid and dashed lines are synthetic RFs of the optimum velocity models with (solid lines) and without
(broken lines) a low-velocity layer in the midcrust in the PHS plate. The optimal velocity models are shown in the left panel.
(b) Same as Figure 9a but for station FUJ.
2. A negative phase after the direct P phases is created by the multiple phases (PsPs or PpSs) in sediment
layers, but the absolute amplitude of phase B is too large to explain by multiple phases alone. The Ps
converted wave at the upper boundary of the low-velocity zone at depths of 13–18 km (Model 1) creates
a large negative amplitude at both stations.
3. The positive amplitudes at about 4–5 s (phase C) at MMS is created by the positive velocity boundary at
20 km for Model 1 and 47 km for Model 2. Phase C at FUJ is created by the positive velocity boundary
at about 26 km for Model 1 and 38–41 km for Model 2.
4. The large negative amplitude at 7–8 s (phase D) is explained only by Model 1, where the PpPs phase
converted at the upper boundary of the low-velocity layer has a negative amplitude at 8 s.
Considering the aforementioned discussion, we conclude that a low-velocity layer at depths of about
13–20 km below MMS and 18–26 km below FUJ, respectively, is required to fit all positive and negative
arrivals of RF around Mount Fuji.
Nakamichi et al. [2007] found a magma chamber of Mount Fuji at depths of 15–25 km from tomographic
imaging of seismic velocity. The low-velocity layer we found in the above analysis may represent the
magma chamber of Mount Fuji. A velocity boundary at depths of 20–30 km below the region where LFEs
of Mount Fuji occurred (Figure 6) probably represents the bottom of this magma chamber.
The depth of conversion can change because we use flat-layered isotropic velocity structures in the grid
search, although polarity reversals that represent dipping or anisotropic layers are found in the observed
transverse components. The depth error by the dipping layers and anisotropic structures is within 5 km as
discussed in section 4.
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a
Table 1. Model Parameters in the RF Grid Search
Bottom Depth (km)
Vp (km/s)
Vs (km/s)
Layer
Name
Lower
Upper
Spacing
Name
Lower
Upper
Spacing
Lower
Upper
Vp/Vs
Model 1
Layer 1
Layer 2
Layer 3
Layer 4
Da
Db
Dc
Dd
1
2
3
4
9
30
40
60
1
1
1
2
v1
v2
v3
v4
1
5
3
5
6
7
6
7
1
1
1
1
0.45
2.89
1.36
2.89
2.73
4.05
2.73
4.05
2.2
1.73
2.2
1.73
Model 2
Layer 1
Layer 2
Layer 3
Da
Db
Dc
1
2
3
9
60
60
1
1
1
v1
v2
v3
1
5
6
6
7
7.5
0.5
0.5
0.5
0.45
2.89
3.47
2.73
4.05
4.34
2.2
1.73
1.73
a
Names, lower bounds, upper bounds, and spacing of model parameters are listed. Vp/Vs values are fixed, and S velocities are calculated using the Vp/Vs ratio in
each layer. Layer 3 in Model 1 represents the low-velocity layer.
5.3. Distinct Positive Phases in the Izu Collision Zone
We find distinct positive velocity boundaries which extend from 40 to 60 km deep below the region where
the IBA has been colliding with central Japan (solid black lines in Figures 5c–5e and 6b). These distinct
phases are consistent with the finding of Asano et al. [1985], who defined a velocity boundary at a depth
of about 40 km below the Izu peninsula from a seismic refraction experiment. Kodaira et al. [2007] and
Takahashi et al. [2009] developed the seismic structure models of the crust and the uppermost mantle
using active source seismic profiling along and across the precollisional IBA, respectively. They found
thickened middle crust below basaltic volcano islands and a 5–10 km thick mafic/ultramafic crust-mantle
transition layer (CMTL) [Takahashi et al., 2009] between the lower crust and the upper mantle along IBA.
They also delineated that the boundary between CMTL and the uppermost mantle is about 30–35 km
deep below volcanic islands. Considering that the middle crust of the Izu collision zone is thicker than
that of volcanic islands because the onshore volume is much larger, the velocity boundary between
CMTL and the uppermost mantle can be deeper than 30–35 km in the Izu collision zone. Combining
previous results with ours, the distinct positive phase we find at a depth of 40–60 km below the Izu
collision zone represents the boundary between CMTL and the uppermost mantle layer.
5.4. The Gap of the Positive Phase Below Mount Fuji
We find a gap of the positive phases that represent the boundary between CMTL and the uppermost
mantle just below Mount Fuji (shown by white arrows in Figures 5d and 6). The gap in the velocity
boundary would suggest that the velocity contrast is locally gradual or that there is no velocity gradient
in the region. To address why this velocity boundary does not continue to just below Mount Fuji, we
stack RFs with a cutoff frequency of 1 Hz and 2 Hz for all the rays that intersect regions A, B, and C in
Figure 10a (same cross section in Figure 6b). The number of seismic rays passing through the target area
is 264 in region A, 163 in region B, and 162 in region C. Stacked RFs with a cutoff frequency of 1 Hz in
Figures 10b and 10d show distinct positive phases at about 5–6 s that represent positive velocity
boundaries between 40 and 50 km, whereas there is a broad negative phase between 4 and 7 s in
Figure 10c. A stacked RF with a cutoff frequency of 2 Hz in region B shows a negative phase at a depth
between 40 and 50 km (gray arrow in Figure 10c) representing that the velocity contrast at about
40–50 km in region B is weaker than that in regions A and C. Furthermore, RFs observed at station MMS
(Figure 4c) show that the distinct positive phases at about 6 s are not observed with the back azimuth
range between 150 and 200°. These support the idea that S velocity contrast between CMTL and the
upper mantle becomes locally weak just below Mount Fuji. The existence of this gap just below the
magma chamber of Mount Fuji also implies that the upwelling magma of Mount Fuji from PAC plate
passes through the area.
5.5. The Seismic Structure Around Mount Fuji
Figure 11 gives a schematic illustration of our result in the N-S and SW-NE cross sections. In the N-S cross
section (Figure 11a), a velocity boundary at depths of 40–60 km below the Izu collision zone represents the
boundary between CMTL and the uppermost mantle. There is a gap of this boundary just below Mount
Fuji, representing a weaker velocity contrast zone at a depth of 50 km. The upper boundary of the
subducting PHS slab at the Izu collision zone is the lower crust [Arai, 2011; Arai et al., 2014]. The positive
KINOSHITA ET AL.
IMAGING MOUNT FUJI BY RECEIVER FUNCTIONS
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Journal of Geophysical Research: Solid Earth
10.1002/2014JB011522
Figure 10. (a) Cross section of an RF amplitude of region 5 in Figure 6b. Black dashed lines indicate regions shown in
Figures 10b–10d. Stacked RFs whose raypaths intersect (b) region A, (c) region B, and (d) region C. (left) Stacked RFs with
a cutoff frequency of 1.0 Hz. (right) Stacked RFs with a cutoff frequency of 2.0 Hz. We plot the arrival time of Ps waves
converted at the depths of 40 km, 50 km, and 60 km derived from Nakamichi et al. [2007]. The gray arrow in Figure 10c
(right) represents the negative phase discussed in section 5.4.
boundary at depths of about 20–30 km is the bottom boundary of the magma chamber of Mount Fuji. In the
SW-NE cross section (Figure 11b), we find the same velocity boundaries where IBA is colliding with central
Japan as seen in the N-S cross section in Figure 11a. We also find the Moho boundary where the back-arc
crust and the fore-arc crust of the PAC plate are subducting. A velocity boundary is at depths of about
50–60 km where the back arc of the PHS is subducting. We interpret this positive phase as the velocity
boundary in the upper mantle of the PHS plate.
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Acknowledgments
The data used in this study are from the
Volcano Research Center database, ERI,
and can be obtained from the Chief of
the Volcano Research Center, Minoru
Takeo, via direct contact ([email protected]). Hypocenter data are
available at the Japan Meteorological
Agency (JMA). We are grateful to Martha
Savage for providing important advice
and comments. Haruhisa Nakamichi
gave us helpful advice and shared his
data for the seismic structure and
low-frequency earthquakes below
Mount Fuji. We also thank Mie Ichihara
for important advice and discussions.
The comments from reviews by two
anonymous reviewers and the Associate
Editor improved the manuscript. We
thank NIED, the Hot Spring Research
Institute of Kanagawa Prefecture, and
Nagoya University for waveform data,
and JMA for waveform data and the
unified hypocenter data. S. K. was
supported by the Grant-in-Aid for
Scientific Research(09J08208) from the
Japan Society for the Promotion of
Science. Many figures were originally
created using the Generic Mapping Tools
[Wessel and Smith, 1991] provided by
Hawaii University.
KINOSHITA ET AL.
10.1002/2014JB011522
Figure 11. Schematic illustrations of the crust and upper mantle structure beneath Mount Fuji along the (a) N-S and
(b) SW-NE cross sections. A velocity boundary at depths of 40–60 km below the Izu collision zone represents the boundary
between CMTL and the uppermost mantle. This distinct boundary becomes locally weak at a depth of 50 km just below Mount
Fuji. The positive boundary at depths of about 20–30 km is the bottom boundary of the magma chamber of Mount Fuji. The
upper boundary of the subducting PHS package at the Izu collision zone is the lower crust [Arai, 2011; Arai et al., 2014]. There is
the Moho boundary where the back-arc crust and the fore-arc crust of the PAC plate are subducting.
6. Conclusions
We apply a receiver function analysis to the teleseismic data recorded by seismic stations around Mount Fuji
to estimate the depth of the velocity boundaries in the subducting PHS slab and gain insight into the
magma-plumbing system of Mount Fuji. Our findings can be summarized as follows.
1. Cross sections of RF amplitudes reveal two distinct velocity boundaries around Mount Fuji at 40–50 km
and 20–30 km, which are interpreted as the upper boundary of the uppermost mantle of IBA and the
velocity discontinuity below the region where low-frequency earthquakes of Mount Fuji occurred.
2. We find a gap of the velocity boundary of IBA at about the 50 km depth just below Mount Fuji, which can
be interpreted as the weaker velocity contrast zone representing piled magma of Mount Fuji ascending
from the PAC plate.
3. Forward modeling of RF shows that a low-velocity zone at 13–26 km depth is required to explain all the
characteristics of obtained RFs around Mount Fuji.
4. We interpret a velocity boundary at 20–30 km below Mount Fuji to be related to the bottom boundary of
the magma chamber of Mount Fuji.
References
Aizawa, K., R. Yoshimura, and N. Oshiman (2004), Splitting of the Philippine Sea Plate and a magma chamber beneath Mt. Fuji, Geophys. Res.
Lett., 31, L09603, doi:10.1029/2004GL019477.
Ammon, C. J. (1991), The isolation of receiver effects from teleseismic P waveforms, Bull. Seismol. Soc. Am., 81, 2504–2510.
IMAGING MOUNT FUJI BY RECEIVER FUNCTIONS
3253
Journal of Geophysical Research: Solid Earth
10.1002/2014JB011522
Arai, R. (2011), Multiple collision and subduction structure of the Izu-Bonin arc revealed by integrated analysis of active and passive source
seismic data, Doctoral thesis, Univ. of Tokyo, Tokyo.
Arai, R., T. Iwasaki, H. Sato, S. Abe, and N. Hirata (2014), Contrasting subduction structures within the Philippine Sea plate: Hydrous oceanic
crust and anhydrous volcanic arc crust, Geochem. Geophys. Geosyst., 15, 1977–1990, doi:10.1002/2014GC005321.
Asano, S., et al. (1985), Crustal structure in the northern part of the Philippine Sea plate as derived from seismic observations of Hatoyama-off
Izu Peninsula explosions, J. Phys. Earth, 33, 173–189.
Cassidy, J. (1992), Numerical experiments in broadband receiver function analysis, Bull. Seismol. Soc. Am., 82, 1453–1474.
Fujii, T. (2007), Magmatology of Fuji Volcano [in Japanese with English abstract], in Fuji Volcano, pp. 233–244, Yamanashi Institute of
Enviromental Sciences, Fuji-Yoshida, Japan.
Igarashi, T., T. Iidaka, and S. Miyabayashi (2011), Crustal structure in the Japanese islands inferred from receiver function analysis [in Japanese
with English abstract], J. Seismol. Soc. Jpn., 63, 139–151.
Julia, J. (2007), Constraining velocity and density contrasts across the crust-mantle boundary with receiver function amplitudes, Geophys.
J. Int., 171, 286–301, doi:10.1111/j.1365-2966.2007.03502.x.
Kaneko, T., A. Yasuda, T. Fujii, and M. Yoshimoto (2010), Crypto-magma chambers beneath Mt. Fuji, J. Volcanol. Geotherm. Res., 193, 161–170,
doi:10.1016/j.jvolgeores.2010.04.002.
Kennett, B. L. N., and N. J. Kerry (1979), Seismic waves in a stratified half space, Geophys. J. R. Astron. Soc., 57, 557–583, doi:10.1111/
j.1365-246X.1979.tb06779.x.
Kennett, B. L. N., E. R. Engdahl, and R. Buland (1995), Constraints on seismic velocities in the Earth from traveltimes, Geophys. J. Int., 122,
108–124, doi:10.1111/j.1365-246X.1995.tb03540.x.
Kodaira, S., T. Sato, N. Takahashi, A. Ito, Y. Tamura, Y. Tatsumi, and Y. Kaneda (2007), Seismological evidence for variable growth of crust along
the Izu intraoceanic arc, J. Geophys. Res., 112, B05104, doi:10.1029/2006JB004593.
Ludwig, W. J., J. E. Nafe, and C. L. Drake (1970), Seismic refraction, in The Sea, vol. 4, edited by A. E. Maxwell, pp. 53–84, Wiley, New York.
Nakajima, J., F. Hirose, and A. Hasegawa (2009), Seismotectonics beneath the Tokyo metropolitan area, Japan: Effect of slab-slab contact and
overlap on seismicity, J. Geophys. Res., 114, B08309, doi:10.1029/2008JB006101.
Nakamichi, H., M. Ukawa, and S. Sakai (2004), Precise hypocenter locations of midcrustal low-frequency earthquakes beneath Mt. Fuji, Japan,
Earth Planets Space, 56, 40–43.
Nakamichi, H., H. Watanabe, and T. Ohminato (2007), Three-dimensional velocity structures of Mount Fuji and the South Fossa Magna,
central Japan, J. Geophys. Res., 112, B03310, doi:10.1029/2005JB004161.
Nakamura, H., H. Iwamori, and J. Kimura (2008), Geochemical evidence for enhanced fluid flux due to overlapping subducting plates, Nat.
Geosci., 1(6), 380–384, doi:10.1038/ngeo200.
Oikawa, J., S. Tanaka, T. Tsutsui, H. Katayama, N. Matsuo, and Y. Nishimura (2004), Controlled source seismic exploration of Fuji volcano in
2003 [in Japanese], Chikyu Mon., 48, 23–34.
Park, J., and V. Levin (2000), Receiver functions from multiple-taper spectral correlation estimates, Bull. Seismol. Soc. Am., 90, 1507–1520,
doi:10.1785/0119990122.
Park, J., C. R. Lindberg, and F. L. I. Vernon (1987), Multitaper spectral analysis of high frequency seismograms, J. Geophys. Res., 92,
12,675–12,684, doi:10.1029/JB092iB12p12675.
Salmon, M. L., T. A. Stern, and M. K. Savage (2011), A major step in the continental Moho and its geodynamic consequences: The Taranaki-Ruapehu
line, New Zealand, Geophys. J. Int., 186, 32–44, doi:10.1111/j.1365-246X.2011.05035.x.
Sato, H., T. Iwasaki, K. Kasahara, E. Kurashimo, T. Ishiyama, R. Arai, and T. Nakayama (2012), Lithospheric structure of the Kanto area deduced
by low-fold seismic reflection profiling and earthquake interferometry analysis [in Japanese], FY 2011 Report of the Special Project for
Earthquake Disaster Mitigation in Tokyo Metropolitan Area, pp. 75–136.
Savage, M. K., J. Park, and H. Todd (2007), Velocity and anisotropy structure at the Hikurangi subduction margin, New Zealand from receiver
functions, Geophys. J. Int., 168, 1034–1050, doi:10.1111/j.1365-246X.2006.03086.x.
Seno, T., S. Stein, and A. E. Gripp (1993), A model for the motion of the Philippine Sea Plate consistent with NUVEL-1 and geological data,
J. Geophys. Res., 98, 17,941–17,948, doi:10.1029/93JB00782.
Seno, T., T. Sakurai, and S. Stein (1996), Can the Okhotsk plate be discriminated from the North American plate?, J. Geophys. Res., 101,
11,305–11,315, doi:10.1029/96JB00532.
Shiomi, K. (2013), New Measurements of Sensor Orientation at NIED Hi-net Stations [in Japanese with English abstract], Report 80, pp. 1–20,
NIED. [Available at http://dil-opac.bosai.go.jp/publication/nied_report/PDF/80/80-1shiomi.pdf.]
Shiomi, K., H. Sato, K. Obara, and M. Ohtake (2004), Configuration of subducting Philippine Sea plate beneath southwest Japan revealed from
receiver function analysis based on the multivariate autoregressive model, J. Geophys. Res., 109, B04308, doi:10.1029/2003JB002774.
Takahashi, M., and K. Saito (1997), Miocene intra-arc bending at an arc-arc collision zone, central Japan, Isl. Arc, 6, 168–182, doi:10.1111/
j.1440-1738.1997.tb00168.x.
Takahashi, N., S. Kodaira, Y. Tatsumi, M. Yamashita, T. Sato, Y. Kaiho, S. Miura, T. No, K. Takizawa, and Y. Kaneda (2009), Structural variations of
arc crusts and rifted margins in the southern Izu-Ogasawara arc-back arc system, Geochem. Geophys. Geosyst., 10, Q09X08, doi:10.1029/
2008GC002146.
Tonegawa, T., K. Hirahara, T. Shibutani, and N. Fujii (2006), Lower slab boundary in the Japan subduction zone, Earth Planet. Sci. Lett., 247,
101–107, doi:10.1016/j.epsl.2006.04.036.
Tsukui, M., M. Sakuyama, T. Koyaguchi, and K. Ozawa (1986), Long-term eruption rates and dimensions of magma reservoirs beneath quaternary
polygenetic volcanoes in Japan, J. Volcanol. Geotherm. Res., 29, 189–202.
Ukawa, M. (2005), Deep low-frequency earthquake swarm in the mid crust beneath Mount Fuji (Japan) in 2000 and 2001, Bull. Volcanol., 68,
47–56, doi:10.1007/s00445-005-0419-5.
Wessel, P., and W. H. F. Smith (1991), Free software helps map and display data, Eos Trans. AGU, 72(41), 445–446, doi:10.1029/90EO00319.
Wirth, E. A., and M. D. Long (2012), Multiple layers of seismic anisotropy and a low-velocity region in the mantle wedge beneath Japan:
Evidence from teleseismic receiver functions, Geochem. Geophys. Geosyst., 13, Q08005, doi:10.1029/2012GC004180.
KINOSHITA ET AL.
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