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Geological Society, London, Special Publications
Magma pathway and its structural controls of Asama
Volcano, Japan
Yosuke Aoki, Minoru Takeo, Takao Ohminato, Yutaka Nagaoka
and Kiwamu Nishida
Geological Society, London, Special Publications 2013, v.380;
p67-84.
doi: 10.1144/SP380.6
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Notes
© The Geological Society of London 2013
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Magma pathway and its structural controls of Asama Volcano, Japan
YOSUKE AOKI1*, MINORU TAKEO1, TAKAO OHMINATO1,
YUTAKA NAGAOKA1,2 & KIWAMU NISHIDA1
1
Earthquake Research Institute, University of Tokyo, 1-1 Yayoi 1,
Bunkyo-ku, Tokyo 113-0032, Japan
2
Present address: Japan Meteorological Agency, 3 – 4 Otemachi 1,
Chiyoda-ku, Tokyo 110-8122, Japan
*Corresponding author (e-mail: [email protected])
Abstract: Asama Volcano, Japan, is one of the most active volcanoes in the Japanese islands.
Recent development of geophysical monitoring in Asama Volcano allows us to infer the magma
pathway and its structural controls beneath the volcano. Combining geodetic data and precise
earthquake locations during recent eruptions suggests that the magma intrudes several kilometres
to the west of the summit to a depth of about 1 km below sea level as a nearly east–west-trending
dyke. The vertically intruded magma then moves horizontally by several kilometres to beneath the
summit before it ascends vertically to make the surface. Combining the P-wave velocity and the
resistivity structure shows that the intrusions are under structural controls. Frozen and fractureless
magma associated with volcanic activity until 24 000 years ago impedes the ascent of rising
magma on its way to the surface. The S-wave velocity structure inferred from ambient noise tomography reveals a low-velocity body beneath the modelled dyke. From independent information,
we have inferred that this low-velocity body is likely to be a magma chamber.
Asama Volcano, in central Japan, is one of the arc
volcanoes associated with subduction of the Pacific
Plate (Fig. 1). The Asama volcanic complex has
been active for about 100 000 years, in which the
centre of activity has migrated from west to east.
The currently active Asama Volcano lies at the
eastern end of this complex and was formed about
24 000 years ago (Aramaki 1963; Aizawa et al.
2008).
The magma of Asama Volcano is mainly andesite, typical of arc volcanoes, so that eruptions are
sometimes explosive. Explosive eruptions with a
Volcano Explosivity Index (VEI; Newhall & Self
1982) of 4 or larger occurred in 1108 AD and 1783
AD. The activity was high between the 1900s and
1960s, with eruptions of VEIs of up to 3. In particular, eruptions in the 1930s were especially explosive, with ballistic volcanic bombs of 1 m or more
in diameter reaching as far as 4 km or more from
the summit (Minakami 1942).
The eruptions have become less explosive during the last few decades but there were moderatesized eruptions with VEIs of 2 in 1982, 1983 and
2004, and, more recently, minor eruptions with
VEIs of 1 in 2008 and 2009. Trace amounts of
ashfall in the Tokyo metropolitan area, 140 km
from Asama Volcano, associated with a minor eruption in February 2009 suggests that the Tokyo
metropolitan area, with a population of more than
20 million, is potentially at risk of volcanic
hazards from the volcano during future larger
eruptions. Table 1 summarizes major and recent
eruptions of Asama Volcano.
With this background, Asama Volcano has
attracted public interest, and has been well documented and recorded. Observations at Asama
started in 1910, motivated by eruptions and a seismic swarm in 1909 (Omori 1912). This was one of
the earliest seismic observations made of an active
volcano. A major achievement during the early
days of seismic observation was the classification
of earthquakes from the observed waveforms;
Minakami (1960) classified the observed seismograms into what we now call volcano– tectonic
earthquakes, low-frequency earthquakes, explosive
earthquakes and harmonic tremors.
After the turn of the twenty-first century, we
started to build a modern monitoring network to
gain more insights into the internal structure of the
volcano and the mechanics of magma transport
beneath the volcano. Although seismic observations
in Asama have been continuing for about 100 years,
there were only six seismometers in 1998 and 11 in
2004. The seismic network has grown rapidly since
then, with currently 30 seismometers as of March
2012, 16 of which are broadband sensors. Continuous GPS observations started in the mid-1990s. As
of 2004, there were nine continuous GPS sites
within 20 km of the summit, two within 4 km and
none within 1 km. The network has grown since
From: Pyle, D. M., Mather, T. A. & Biggs, J. (eds) 2013. Remote Sensing of Volcanoes and Volcanic Processes:
Integrating Observation and Modelling. Geological Society, London, Special Publications, 380, 67– 84.
First published online March 4, 2013, http://dx.doi.org/10.1144/SP380.6
# The Geological Society of London 2013. Publishing disclaimer: www.geolsoc.org.uk/pub_ethics
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68
Y. AOKI ET AL.
also have 10 tiltmeters (Fig. 2c), nine microphones
(Fig. 2d) and two cosmic ray muon detectors
(Fig. 2c), as of March 2012.
Overview of recent unrest
Seismic and geodetic observations show that the
activity of Asama Volcano is not steady over time.
Indeed, the seismic activity of the volcano tends to
be higher when the volcano inflates and lower
when the volcano deflates. This section describes
the overview of the seismic and geodetic observations associated with unrest during the era of
modern instrumentation since the mid-1990s and
eruptions during that period (2004, 2008 and 2009).
While Synthetic Aperture Radar (SAR) interferometry often plays an important role in monitoring volcano deformation (e.g. Dzurisin 2007), they
make little impact in monitoring Asama Volcano
for some reason. The area around Asama Volcano
is densely vegetated, as in other regions in the Japanese islands, so that only L-band SAR data works
for interferometric analyses. Also, snow coverage
on Asama between December and early May prevents us from retrieving meaningful displacement
signals associated with unrest.
Fig. 1. Tectonic setting around Asama Volcano. Plate
boundary and active volcanoes are denoted by solid lines
and black rectangles, respectively. Asama Volcano is
shown by a red triangle. PAC, PHS, EUR and OKH stand
for the Pacific, Philippine Sea, Eurasia and Okhotsk
plates, respectively. The converging velocity of PAC
with respect to OKH derived by MORVEL2010 (DeMets
et al. 2010) is also shown.
then and currently we have 15 continuous GPS
sites within 20 km of the summit, four within
4 km and two on the rim of the summit crater. We
Unrest before 2004
Figure 3 compares the temporal variations in the
monthly number of volcano –tectonic earthquakes
and that of all volcanic earthquakes with changes
in GPS baseline length for sites 0221 and 0268
(see Fig. 2b for the location) between 2000 and
2011. The baseline extensions between these two
GPS sites indicate magma intrusions beneath the
western flank of Mt Asama because the baseline
crosses the inferred dyking area, which is deduced
from geodetic data during 2004 and 2008– 2009
Table 1. Summary of major and recent eruptions in Asama Volcano. Dates, VEI, maximum
column heights from the summit, ejected volumes and SiO2 content of the ejecta are shown
Year
Dates
VEI
Max. column
height (m)
Ejected
volume (m3)
SiO2
content (%)
1108
1783
1973
1982
1983
2003
2004
2008
2009
29 August– 11 October
8 May – 15 August
1 February – 24 May
26 April, 2 October
8 April
6 February – 18 April
1 September – 9 December
10 – 14 August
2 February – 27 May
5
4
2
2
2
1
2
1
1
N/A
N/A
4000d
1500g
1400g
500g
3500 – 5500 f
400
1000g
6.2 × 108a
5.1 × 108c
4.4 × 105e
3.3 × 105e,h
59 – 61b
60 – 63c
60 – 61f
N/A
N/A
N/A
61 – 76k
N/A
66 – 73l
a
3.8 × 102j
6.0 × 104j
Trace
1.0 × 104l
Miyahara (1991); bAramaki (1963); cYasui & Koyaguchi (2004); dMiyazaki (2003); eHayakawa (1999);
Nakada et al. (2005); gJapan Meteorological Agency pers. comm; hTotal volume associated with the 1982 and
1983 eruptions; iYoshimoto et al. (2005); jShimano et al. (2005); kMaeno et al. (2010).
f
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MONITORING ASAMA VOLCANO
138°25'
69
138°35'
138°30'
36°30'
(a) Seismometers
138°20' 138°25' 138°30' 138°35' 138°40'
(b) GPS
0221
36°30'
TASH
36°25'
KAHG
0268
AVOG
KVCO
36°25'
36°20'
km
0
km
5
36°15'
36°20'
0
5
36°28'
36°28'
(c) Tiltmeters+Muon
(d) Microphones
36°26'
36°26'
36°24'
36°24'
km
km
0 1 2
0 1 2
36°22'
36°22'
138°28' 138°30' 138°32' 138°34' 138°36'
138°28' 138°30' 138°32' 138°34' 138°36'
Fig. 2. Spatial distribution of: (a) seismometers in circles; (b) GPS sites in diamonds; (c) tiltmeters in rectangles and
cosmic ray muon detectors in stars; and (d) microphones in triangles. In all panels, the location of the summit is denoted
by green triangle. The location of Mt Kurofu is denoted by a red rectangle in (a).
unrest (Aoki et al. 2005; Takeo et al. 2006). The
hypocentre distribution before 2004 demonstrates
that the volcano –tectonic earthquakes occurred
under the western side of Mt Asama and the other
volcanic earthquakes occurred beneath the summit
of Mt Asama, although their precision was not as
good as that after 2004 owing to the lack of dense
seismic network. Because the inferred dyke locations associated with subsequent unrest in 2004,
2008 and 2009 are also to the west of the summit
(Aoki et al. 2005; Murakami 2005; Takeo et al.
2006), it seems reasonable to assume that the overall trend of hypocentral distribution and the dyke
location had not changed during the last decade.
Figure 3 shows that the number of volcano– tectonic
earthquakes increased twice from October 2000
to April 2001 and from May 2002 to August 2002
before the 2004 eruption. Both seismic activations
coincide with the extensions of the GPS baselines,
suggesting that the volcano –tectonic earthquakes
were associated with the intrusion of magma
beneath the western flank of Mt Asama.
The contractions of the GPS baseline were
observed three times: prior to March 2000; from
July 2001 to February 2002; and from March 2003
to April 2004. Because the observed contraction
is too large to be explained by tectonic motion of
non-volcanic origin, we would speculate that it
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Y. AOKI ET AL.
400
(a)
200
0
2000
# of VT EQs
# of volcanic EQs
70
2002
2004
2006
2007
2008
2010
2011
2012
2001
2002
2004
2006
2007
2008
2010
2011
2012
2010
2011
2012
15
10
(b)
5
0
2000
60
Baseline length (mm)
2001
(c)
40
0221–0268
20
0
–20
–40
–60
2000
KVCO–TASH
2001
2002
2003
2004
2005
2006
Year
2007
2008
2009
Fig. 3. Temporal evolution of seismic activity and ground deformation between 2000 and 2011. Eruptive periods are
shown by grey shades. (a) Daily counts of volcanic earthquakes with unclear P and S arrivals triggered by the data
acquisition system. (b) Daily counts of volcano–tectonic earthquakes triggered by the data acquisition system.
(c) Relative changes in baseline length between 0221 and 0268 (dots) and KVCO and TASH (crosses), respectively.
Location of the stations are denoted in Figure 2. Note that the distance between 0221 and 0268 increased significantly
due to the 2011 Tohoku-oki earthquake. However, the distance between KVCO and TASH did not change much because
KVCO–TASH strikes almost north–south while the 2011 Tohoku-oki earthquake generated mostly east– west
expansion without any significant north–south expansion or contraction. Outliers in late 2008 in the KVCO– TASH
baseline are due to instrumental problems.
may indicate migrations of intrusive magma from
under the western flank of Mt Asama. While the
exact direction of the magma migration is unknown
because the change in GPS baseline length between
950221 and 950267 is only sensitive to the inflation and deflation under the western flank of Mt
Asama, the magma may have either migrated vertically down to depth or eastwards towards the
vent at the time of the baseline contraction. The
seismicity was relatively low during the first contraction period except for the last several months,
while it was relatively high during the latter two
contracting periods (Fig. 3). The maximum temperature of the crater had exceeded 200 8C since
autumn 2002 (Japan Meteorological Agency 2005)
with tiny eruption in February 2003, suggesting
that the shallow part of the vent had been at an
elevated temperature state from the middle of
2002. These observations suggest the essential differences between the first and later contractions.
2004 eruption
The 2004 eruption, the first moderate-sized eruption since 1983, was preceded by elevated thermal
activity and increased emission of SO2 from about
2000, with a few minor eruptions between February and April 2003 (Nakada et al. 2005; Takeo
et al. 2006). While seismic activity was high in
2004, it did not increase towards the 2004 eruption
(Fig. 3). However, continuous GPS data shows
that the north–south extension of the volcano
started in late July 2004, about 5 weeks before the
first eruption of 2004 (Fig. 3).
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MONITORING ASAMA VOLCANO
71
2004/01/01~2011/09/27
138°30'
138°32'
(a)
(b)
km
0
1
2
36°26'
–2000
–1000
0
1000
(c)
Jan. Jan. Jan. Jan. Jan. Jan. Jan. Jan.
2004
2006
2008
2010
2000
36°24'
–3 –2 –1 0 1 2 3
depth [km]
N=1001
Fig. 4. Distribution of relocated hypocentres around Asama Volcano between January 2004 and September 2011. Each
hypocentre is coloured by its depth. (a) Distribution of epicentres with the approximate location of dyke intrusions
during crises (red line). (b) North–south cross-section of hypocentres. (c) Temporal evolution of hypocentres. The
horizontal axis represents the longitude of the hypocentres.
The 2004 eruption started on 1 September 2004
as a Vulcanian eruption with a VEI of 2, with the
ash plume reaching 3.5–5.5 km above the summit
or 6–8 km above sea level. After a short period of
quiescence, Strombolian eruptions started on 14
and 16–18 September before Vulcanian eruptions
occurred on 23 and 29 September, 10 October, 14
November and 9 December. The total amount
of ejected tephra is estimated to have been about
1.6 × 108 kg (Nakada et al. 2005). The magma
involved in the 2004 eruption was andesite, with a
composition similar to magmas erupted over the
past 10 000 years.
A lava dome emerged at the summit by 16 September, when continuous Strombolian eruptions
started (Nakada et al. 2005). Combining this with
the change in magma composition of increasing
amounts of juvenile magma in the eruptive products
by that time suggests that juvenile magma may
have ascended to the summit crater by that time.
The lava dome deflated afterwards at its centre, indicating that at least a part of the ascended magma
had descended to depth. The size of the lava dome,
derived from SAR observations, is estimated to be
about 2 × 106 m3 or 5 × 109 kg at its maximum, if
a rock density of 2500 kg m23 is assumed (Nakada et al. 2005; Oki et al. 2005). This value is larger
than the amount of ejected tephra by more than
an order of magnitude (Yoshimoto et al. 2005),
indicating that only a fraction of magma involved
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72
Y. AOKI ET AL.
in the volcanic activity had been ejected, as universally seen in active volcanoes.
Takeo et al. (2006) precisely relocated hypocentres of more than 500 volcanic earthquakes
between 1 January 2004 and 19 October 2005, spanning the period of the 2004 unrest, using a doubledifference algorithm (Waldhauser & Ellsworth
2000). The relocated seismicity reveals a sharp
image composed of two groups (Fig. 4). One forms
a WNW–ESE-trending zone at a depth ranging
between 1 and 1.5 km below sea level. The eastern
end of this seismic zone lies beneath the summit
crater and extends westward horizontally by about
2 km. The other group forms a narrow vertical
seismic zone extending from the eastern edge of
the other group to just under the summit crater.
We also modelled the ground deformation field
between June 2004 and March 2005 by inverting
for length, width, depth, dip angle, strike direction,
location and amount of opening of an intruded
rectangular dyke in an elastic, homogeneous and
isotropic medium (Okada 1985). While topography
or material heterogeneity plays a role in the deformation field (Du et al. 1997; Williams & Wadge
2000), we neglected them in the present study to
make the inverse problem more tractable. Neither
a spherical source nor a fault dislocation model
explains the ground deformation pattern. Instead,
the observations are fitted with a dyke intrusion to
the western flank of Mt Asama (Fig. 5). The intruded
volume was estimated to be 6.8 × 106 m3 (Aoki
et al. 2005; Takeo et al. 2006), which is about three
times larger than the volume of emitted magma
during the eruption (2 × 106 m3: Nakada et al.
2005). While the observed deformation is not
large enough to constrain well all of the model
parameters, the horizontal location (+2 km with
a 95% confidence) and the total volume of the
intruded dyke (+30% with a 95% confidence) are
relatively well constrained (Aoki et al. 2005;
Takeo et al. 2006). The eastern part of the modelled dyke overlaps with the western end of the
relocated hypocentre distribution (Fig. 4). The
depth at the top of the dyke, whose standard deviation is estimated to be 1.3 km, coincides with the
depth range of the volcano –tectonic earthquakes.
The distribution of dyke-induced seismicity reflects
the distribution of ambient stresses that are near
failure; thus, the seismicity might be much more
concentrated near the dyke tip (Rubin & Gillard
1998). The eastern end of horizontally elongated
seismicity is connected with the narrow vertical
seismic zone of volcanic earthquakes extending
from 1 km below sea level to just under the eruptive
summit crater (Fig. 4). We interpret that this distribution of hypocentres represents the magma pathway beneath Mt Asama.
A sudden extension of the GPS baseline length
between sites 950221 and MIYO (Fig. 2b) was
detected between 21 and 22 July 2004, suggesting
Fig. 5. Comparison between the observed displacements between June 2004 and March 2005, and the calculated
displacements from a dyke to the west of the summit shown by a red line. Ellipses represent 1s uncertainties of the
observed displacements. A green triangle indicates the summit of Asama Volcano.
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MONITORING ASAMA VOLCANO
a nearly vertical magma intrusion to the western
flank of Mt Asama (Aoki et al. 2005; Murakami
2005; Takeo et al. 2006). In addition, volcanic
glows had begun to be observed during the last 10
days of July 2004, and the maximum temperature
at the bottom of the crater exceeded 500 8C after
the sudden extension of the baseline length (Japan
Meteorological Agency 2005). These observations
indicate that the elevated temperature at the shallow part of the vent resulted from the magma
intrusion.
Integrating the observations discussed above,
the two baseline contractions after 2000 (Fig. 3)
appear to be caused by the migration of the intrusive magma from the dyke beneath the western
flank to the vent of Mt Asama. The gradual supply
of magma into the vent had induced a gradual activation of volcanic earthquakes from the middle
of 2001 to the first eruption on 1 September 2004.
The relocated hypocentre distribution in Figure 4
shows that almost all earthquakes from January to
1 September 2004, when the 2004 eruption started,
lie in the shallower part of the vent, with a depth
shallower than 1 km above sea level. These suggest that the magma had ascended gradually to
about 1 km above sea level by at least June 2004,
leading to the increased internal pressure at the shallowest part of the conduit. It then resulted in an
enhanced seismic activity at the shallower part
73
of the vent. Before the eruption on 1 September,
magma migration from the deep chamber to the
shallower portion of the vent was relatively slow
probably due to a plug at the top of the vent. After
the eruption, the pressure release by removal of the
plug may have made the magma migration much
easier than before, resulting in more volcano –
tectonic earthquakes. This speculation is supported by the fact that partially molten country rocks
(rhyolite tuff) are found among the 2004 eruption
products (Nakada et al. 2005).
2008 – 2009 eruption
After 4 years of repose, the seismic activity started
to elevate in July 2008, 1 month before the first eruption. The earthquakes are mainly located approximately 1 km to the west of the summit at about
1 km below sea level (Fig. 4). The seismic activity
further increased on 8 August, leading to eruptions
on 10, 11 and 14 August, which were minor with
ash plumes reaching a height of 200– 400 m above
the summit. The eruptive products did not include
juvenile materials, leading us to infer that the 2008
eruption is actually the ejection of the remnant
of magma associated with the 2004 eruption.
Since September 2008, the elevated seismic
activity has been observed approximately 2 km to
Fig. 6. Images of the summit taken on 9 September 2008 (left) and 2 February 2009 (right) from a surveillance camera
located on the western rim of the summit. The left-hand panels show that the bottom of the summit crater is located in the
lower centre of the image. The right-hand panels indicate that the eruption occurred from the left-hand side of the image.
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74
Y. AOKI ET AL.
Fig. 7. Images of the summit crater taken from a surveillance camera located on the western rim of the summit. These
images were taken on 16 May 2009 (top) compared with the one taken on 9 September 2008 (bottom). The most
prominent difference between the two images is shown by white ellipses, indicating the formation of a new vent by May
2009. This vent was probably formed during the February 2009 eruption. Otherwise, there are few morphological
changes between the two images, implying that the February 2009 eruption took place at the single vent.
the west of the summit at about 1 km below sea level
(Fig. 4). Subsequently, the trace amount of ash emission was identified in January 2009. On 1 February,
a day before the largest eruption in 2009, seismicity
right beneath the summit increased and a tiltmeter
approximately 3 km from the summit tilted away
from the volcano showed a signal of volcano inflation, leading to a summit eruption the next day.
Subsequently, minor eruptions from the summit
occurred intermittently on 9–16 February, 15
March, 14 and 30 April, and 3 and 27 May, with a
plume height of less than 500 m.
The largest eruption during the 2008–2009 crisis
was that on 2 February but it was smaller than that
on 1 September 2004. The ejected mass due to
the eruption on 2 February was 2.0– 2.4 × 107 kg,
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MONITORING ASAMA VOLCANO
approximately 20% of that associated with the 1
September 2004 eruption (Maeno et al. 2010). The
ejected materials included only trace amounts of
juvenile materials, suggesting that the eruption was
merely an ejection of near-surface materials rather
than a result of magma migration from depth
(Maeno et al. 2010). This is consistent with insignificant density changes observed by muon radiography (Tanaka et al. 2009).
A surveillance camera on the western rim of the
summit recorded the 2 February eruption. Figure 6
indicates that the eruption occurred on the left-hand
side of the image, which is translated as the NW side
of the crater. Note that the camera points to the
centre of the crater, where the temperature is relatively higher than the surroundings during the dormancy (Fig. 6).
Visual inspection of the crater from the summit
confirms the eruption from the NW part of the
summit. Comparison between photographs taken
before and after the eruption from the same location (Fig. 7) shows that a new vent, with a radius
of 10 m or so, is formed by the 2 February eruption.
It also shows that no prominent changes occurred
75
otherwise, indicating that the eruption occurred
from a single vent. Figure 7 shows that the vent
opens westwards, consistent with the observation
that ballistic materials associated with the eruption
distributed exclusively to the west of the summit
(Fig. 6b– d).
Surface deformation associated with this unrest
has also been monitored by continuous GPS sites.
The area around the volcano was under compression before the summer of 2008, as is usually
observed in volcanoes during dormancy (Murakami
2005; Takeo et al. 2006). The western side of
the volcano started to exhibit north–south extension
from July 2007, a month before the first eruption in
2008 (Fig. 3). The volcano then expanded afterwards quasi-linearly until June 2009. This extension
was due to dyke intrusion to the west of the summit,
as in the 2004 case. In addition to this, there was
expansion near the immediate vicinity of the summit. KAHG (Fig. 2), located on the eastern rim
of the summit, displaced eastsoutheastwards by
approximately 25 mm in the latter part of 2008,
although missing observations prevent us from
identifying the time of the onset (Fig. 8a). KAHG
(b) AVOG
20
20
East (mm)
30
10
0
−10
10
0
−10
−20
−30
−30
30
30
20
20
North (mm)
−20
10
0
−10
10
0
−10
−20
−20
−30
−30
40
40
20
20
Up (mm)
Up (mm)
North (mm)
East (mm)
(a) KAHG
30
0
−20
−40
1/08
0
−20
4
7
10
1/09
4
7
−40
1/08
4
7
10
1/09
4
7
Fig. 8. An example of time series of (a) KAHG and (b) AVOG with respect to a GPS site far enough from Asama
Volcano to be affected by volcanic activity. The location of these sites is shown in Figure 2b.
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76
Y. AOKI ET AL.
then moved northwestwards by about 10 mm in
early 2009. Note that the next closest site, AVOG
(Fig. 2), did not exhibit significant deformation
(Fig. 8b), implying that the source responsible for
the summit deformation is shallow and that the
deformation is localized around the summit.
The displacement field when the summit expansion and the expansion to the west was ongoing
(June–December 2008) is fitted well by an east –
west-striking dyke with its top at 0.7 km below
sea level and a shallow source responsible for the
summit inflation (Fig. 9). The east –west-striking
dyke is located near the one associated with the
2004 unrest, indicating that the dyke intrusion to
the west of the summit is ubiquitous during unrest
of Asama Volcano. This also implies an established magma pathway to the west of the summit.
The volume of intrusion is estimated to be 1.64 ×
106 m3, approximately 25% of that during the
2004 unrest.
Because there is only one GPS site on the
summit, it is impossible to constrain the shape of
the source responsible for the summit deformation
even if the source is assume to be located right
beneath the summit. The expansion of a spherical source can be ruled out because the pressure
change well above the rock fracture toughness is
138°20'
138°25'
required to fit the observed horizontal displacements, which are about 25 mm at 250 m from the
summit and almost zero (at least less than 5 mm)
at 4 km from the summit. For similar reasons, we
can rule out the pressurization of the conduit, with
its ends either open or closed, from viable mechanisms. Traction to the conduit with a radius of,
say, 50 m seems to be a viable mechanism to be
consistent with the observed displacements and the
fracture toughness of rocks. If we assume a conduit with a radius of 50 m extending from the summit and 1 km below sea level, where the vertically
elongated seismicity marks its bottom (Fig. 4), the
required traction on the conduit wall is about
15 MPa, which seems to be reasonable considering
a confining pressure at a depth of 1.5 km below
the surface, for example, is about 30 MPa if an
average rock density of 2000 kg m23 is assumed.
Active source seismic tomography
As seen in the previous section, precise earthquake
locations and deformation fields enabled us to delineate a rather sinuous magma pathway during unrests
of Asama Volcano (Takeo et al. 2006). This raises
two questions. What makes the magma pathway
138°35'
138°30'
138°40'
5 mm
36°30'
36°30'
KAHG
36°25'
36°25'
36°20'
36°20'
km
0
5
36°15'
10
36°15'
138°20'
138°25'
138°30'
138°35'
138°40'
Fig. 9. Comparison between the observed displacements between June and December 2008 (black) when both the
western part of the volcano and the shallower depths beneath the summit are inflating and the calculated displacements
with the optimum dyke (red) that fits the observed displacements, excluding those of KAHG which are clearly affected
by inflation of a shallower origin.
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MONITORING ASAMA VOLCANO
winding? If and how is the magma pathway structurally controlled? To address these questions, we conducted an active source seismic experiment in
October 2006 with a dense deployment of temporary seismometers (Aoki et al. 2009a, b). Active
source seismic experiments can take advantage of
known locations and origin time of artificial sources.
The seismic exploration was conducted with
five dynamite sources of 250 –300 kg and 464 temporary 2 Hz seismometers (Mark Products L22D),
with an average spacing of approximately 150 m
(Fig. 10). To focus on seismic velocity structures around the area of dyke intrusion during the
2004 eruption, we constructed two main lines of
seismometers striking roughly north–south and
east –west, respectively, crossing the area of dyke
intrusion (Fig. 10).
The observed waveforms indicate that seismic
waves that do not traverse the area of dyking have
the maximum energy at first arrivals, while those
traversing the area of dyking have different waveforms with more energy in later arrivals (Aoki
et al. 2009b). This feature is probably due to the
scattering of seismic waves at the area of dyking
where the seismic structure is likely to be quite inhomogeneous. We also note that seismic waves traversing the summit are subject to substantial scattering.
We then constructed P-wave velocity models to
be consistent with first arrivals in each trace first
Fig. 10. Spatial distribution of active sources (red stars)
and seismometers (black dots). The summit of Asmaa
Volcano is shown by a green triangle. Bouguer gravity
anomaly with a reduction density of 2670 kg m23
(Geological Survey of Japan 2000). The approximate
location of the dyke intrusion during crises is marked
in yellow.
77
by forward modelling (Zelt & Smith 1992) and
then applying a regularized inversion of the travel
times with the initial model obtained through the
previous forward modelling (Zelt & Barton 1998).
Figure 11 shows the P-wave velocity structure
for both profiles, in which Figure 11a, b depicts
the comparison between observed and calculated
travel times and P-wave velocity structure for the
north –south profile, respectively, and Figure 11c, d
depicts those for the east– west profile, respectively. Figure 11a, d indicates that the calculated travel times fit well with the observed travel times for
both profiles.
Figure 11b clearly shows a high-velocity zone
around S3 (distance 0 km), at a depth of 2 km
below sea level and shallower, the area of dyke
intrusion during the 2004 eruptions inferred from
seismic and geodetic measurements (Takeo et al.
2006). Figure 11d shows that the velocity to the
west of the summit is faster than that to the east of
the summit. Combining this with the velocity along
the north–south profile suggests that the highvelocity zone strikes east –west, with its top deepening to the east. The north–south extent of the
high-velocity zone is as narrow as approximately
5 km. We conclude that this is due to solidified
magma resulting from repeated dyke intrusions
because the location of the high-velocity zone
roughly coincides with the area of dyking associated with the 2004 eruptions. Note that the highvelocity zone is not formed by a single dyking, for
example, in 2004. Geodetic data show that the
dyke intruded in 2004 is less than 1 m thick
(Takeo et al. 2006), too thin to be imaged in this
analysis if the high-velocity zone was formed by a
single dyking in 2004.
Active and passive seismic surveys found highvelocity zones in many active volcanoes (e.g.
Zollo et al. 1996, 1998; Okubo et al. 1997; Tanaka
et al. 2002; Yamawaki et al. 2004). We interpret the
high-velocity zone as solidified magma due to past
intrusions, as previous studies did. Here we would
like to note that the high-velocity zone we found
is consistent with the regional geology that Quaternary volcanoes distribute roughly east –west, the
eastern end and the youngest of which is Asama
(Aramaki 1963). Our result is also consistent with
insights gained from other geophysical observations; that is, precise earthquake locations, ground
deformation, gravity anomaly and electromagnetic
structure.
Figure 10 indicates that the spatial distribution
of the Bouguer gravity in the area has east –westtrending local maxima to the west of the summit,
the area of high velocity, and the eastern end of
which is at Asama. This saddle-shaped gravity
high suggests that the high-velocity zone is imaged
by a high-density zone, consistent with our
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78
Y. AOKI ET AL.
Fig. 11. P-wave velocity structure obtained by travel time inversion. (a) Comparison of observed and calculated travel
times for the north– south profile. (b) P-wave velocity structure along the north–south profile. (c) Same as (a) but for the
east–west profile. (d) P-wave velocity structure along the east– west profile.
Fig. 12. Spatial distribution of seismic sites (black dots) used in the ambient noise tomography. Three regions for
phase-velocity measurements are circled. The location of Asama Volcano is also shown. Geographical location of the
area is shown in the inset.
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MONITORING ASAMA VOLCANO
hypothesis that the high-velocity zone is due to the
solidification of repeatedly intruded magma. Note
that when magma is cooled slowly, it will become
fractureless, thus dense, rock with high seismic
velocity.
Electromagnetic data show that the area of the
high-velocity zone is less conductive than surrounding areas (Aizawa et al. 2008), also consistent with
our hypothesis because slowly solidified magma
has low permeability, resulting in high resistivity.
Aizawa et al. (2008) referred to this high resistivity,
extending up to about 1 km above sea level, as solidified magma associated with eruptions of Kurofu,
an active centre approximately 3 km to the west of
the current vent, up until 24 000 years ago (Aramaki
1963).
Our hypothesis that the high-velocity zone is due
to solidified magma implies that some amount of
magma should fail to reach the surface during
eruptions. In fact, geodetic data, field observations
and the volume of the lava dome at the summit
measured by Airborne Synthetic Aperture Radar
observations all suggest that in the 2004 eruptions,
79
only a small fraction of intruded magma was
ejected (Oki et al. 2005). We speculate that in
Asama Volcano, much of the intruded magma did
not make the surface but was arrested and solidified in past eruptions, as well as in the 2004
eruptions.
Ambient noise seismic tomography
While the active source seismic experiment is able
to resolve shallow seismic structure, as shown in
the previous section, it does not resolve any structures deeper than 3 –4 km below sea level (Aoki
et al. 2009b). However, a regional seismic tomography study of the area images broader-scale structures from the lower crust to the upper mantle
(Nakajima & Hasegawa 2007) but is not capable
of imaging the upper crust in detail, where magma
chambers are likely to exist. Here we fill the gap
to image the seismic structure of the upper crust
beneath Mt Asama using seismic ambient noise
(Nagaoka et al. 2012).
Fig. 13. Perturbations in travel times with respect to the reference dispersions and the obtained phase velocities. Top
panels represent the travel time perturbations with respect to those calculated from the reference dispersions at: (a) 0.1–
0.2 Hz, (b) 0.15–0.3 Hz and (c) 0.2– 0.4 Hz, respectively. Bottom panels represent phase velocities at: (a) 0.1–0.2 Hz,
(b) 0.15–0.3 Hz and (c) 0.2– 0.4 Hz, respectively.
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80
Y. AOKI ET AL.
We first divided the whole region into three subregions (Fig. 12) according to a priori knowledge
from a previous study (Nishida et al. 2008), which
shows a clear velocity contrast with slower NW
and faster SE. We then took cross-correlations of
all possible pairs within each subregion, with appropriate pre-processing including mean removal,
detrend and spectral whitening, as done by previous
studies (Bensen et al. 2007).
The obtained cross-correlations show the propagation of the Rayleigh wave trains in all regions
but with different speed. The main wave packet
travels at about 3 km s21 in region 2, much faster
than in other regions where it is about 2 km s21
(Nagaoka et al. 2012), consistent with previous
studies (Nishida et al. 2008). In each region, reference dispersion curves were measured with an
assumption that the seismic structure varies only
with depth within each region.
Once the reference phase velocity is obtained,
the travel time anomaly is measured as a perturbation from the reference phase velocity for all
available station pairs (Fig. 13a –c). Figure 13d
shows that the obtained phase velocity of a frequency range of 0.1–0.2 Hz is slower to the west
of the volcano by up to 10%. Furthermore, the
measured phase velocities between station pairs
were inverted for two-dimensional phase velocity
maps. We applied an iterative non-linear inversion (Rawlinson & Sambridge 2003) to estimate
phase velocities of a given frequency range at grid
points with a spacing of 0.038. Taking the phase
velocity map from a previous study (Nishida et al.
2008) as a initial model, we iterated the inversion
10 times to get the final model. The spatial variations of the Rayleigh wave-phase velocity at a frequency of 0.1–0.2 Hz show a negative velocity
anomaly of up to about 20% approximately 10 km
to the west of the summit (Fig. 13d).
A spike resolution test demonstrates that our data
have enough power to resolve the low velocity of
this size (Nagaoka et al. 2012). While the mapped
low-phase velocity at 0.1–0.2 Hz has a radius of
about 5 km (Fig. 13d), the actual size would be
Fig. 14. S-wave velocity around Asama Volcano. Upper panels represent the plan view at depths of 0 (left) and 5 km
(right) below sea level, respectively. Lower panels represent the cross-sections along the A– B (left) and C–D (right)
baselines, respectively, shown in the upper panels.
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MONITORING ASAMA VOLCANO
smaller because the analysis in this study does not
take the effect of finite frequency into account.
Under an assumption of infinite frequency, which
is employed here, geometrical rays have a resolution to the structure only along the raypath. In
reality, however, rays with finite frequency have a
resolving power to the structure of approximately
one third of the first Fresnel zone width (Yoshizawa
& Kennett 2002), which is approximately 7 km for
the 0.1–0.2 Hz case. Because the actual velocity
anomaly off the geometrical ray but within one
third of the Fresnel zone will be mapped along the
geometrical ray, the mapped velocity anomaly will
be broader than the real one. We thus conclude
that the radius of the velocity anomaly would actually be smaller than 5 km.
Figure 14 shows the three-dimensional S-wave
tomographical model from a collection of the local
one-dimensional S-wave velocity structures. This
figure indicates low S-wave velocity at depths of
5–10 km. Such a low-velocity region to the west of
the summit disappears at higher frequencies (Fig.
13e– f ). Instead, Figure 14 displays high S-wave
velocity anomalies at shallower depths (,3 km)
around the repeated dyking region inferred from
an active source seismic exploration (Aoki et al.
2009b). The S-wave velocity structure is consistent
with the P-wave velocity derived from an active
source seismic experiment (Aoki et al. 2009b). In
addition, the inferred S-wave velocity at the depth
of sea level shows a slower phase velocity to the
east of the summit than to the west, which is also
consistent with the result from an active source
seismic experiment (Aoki et al. 2009b).
81
What makes the low S-wave velocity change
from 5 to 10 km? Here we suggest, with some evidence, that it is a magma chamber. We observed
tilt motions during the February 2009 eruption
(Matsuzawa et al. 2009; Miyamura et al. 2009),
many of which were obtained from the horizontal record of broadband seismometers by taking
advantage of the capability of broadband seismometers in recording tilt motions (Rodgers 1968;
Graizer 2006; Maeda et al. 2011). We excluded
tilt records from sites near the summit, with tilting
up to 7 mrad towards the summit, because they
reflect the mass loss at the summit associated with
the eruption, which is outside of our scope. Other
sites generally tilt towards to the west of the summit by up to 1.1 mrad (Fig. 15), implying a source
offset to the west. We modelled the observed deformation field by a deflating dyke embedded in a
homogeneous, elastic and isotropic half-space
(Okada 1985). The optimum dyke is at its top at
about 1 km below sea level with a NW–SE strike
(Fig. 15), the location of which is right above the
low S-wave velocity anomaly.
The existence of a dyke participating in an eruption suggests that the low S-wave velocity anomaly
represents the magma chamber beneath the dyke
deflating during the 2009 eruption. The crustal
magma chamber has not been imaged at all (Aoki
et al. 2009b) or only as a broad low-velocity zone
Kurofu
Summit
East
–2
0
High
resistivity
High
resistivity
Depth (km)
High Vp
2
Magma pathway
Dike intrusion
4
6
Low Vs
Magma chamber
8
10
Fig. 15. Modelling of tilt motions associated with the 2
February 2009 eruption. The optimum location of the
closing dyke is shown in blue. Black and pink vectors
represent observed and calculated tilting directions,
respectively. Summit of Asama Volcano is shown by a
red triangle.
Fig. 16. Schematics of the magma pathway integrating
seismic and geodetic observations. Black dots represent
seismicity. Magma starts to ascend from the magma
chamber as a dyke. The intruded magma is blocked by a
high compressional wave velocity (Vp) (Aoki et al.
2009a, b) and high resistivity (Aizawa et al. 2008) body
and follows a path represented by the seismicity. The
location of Kurofu is shown in Figure 2a.
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82
Y. AOKI ET AL.
(Lees 2007) by conventional seismic imaging
studies because of technical difficulties. So this
study is the first to delineate the size of a crustal
magma chamber. Also, it has not been inferred
from geodetic data (Aoki et al. 2005; Takeo et al.
2006) because the current distribution of continuous GPS sites is not capable of detecting displacements owing to the inflation or deflation of the
deep chamber. Our results also demonstrate that
ambient noise cross-correlations are capable of
imaging an anomaly with a radius of approximately
5 km at a depth of 5 km or deeper with an optimal
station coverage.
Synthesis
Taken together with discussions above, in Figure 16
we show a synthesis of the magma pathway in the
upper crust beneath Asama Volcano based on our
observations and discussion. When magma rises
up to the upper crust, it is stored in the magma
chamber at a depth of 5–10 km. The magma
chamber is offset by approximately 8 km to the
west from the summit and is represented by a lowvelocity region. Then the magma further moves
up to a depth of about 1 km below sea level
as a dyke. The shallow seismic and resistivity structures (Aizawa et al. 2008; Aoki et al. 2009a, b)
suggest that the magma pathway is subject to structural controls to make a winding path to reach the
surface. Combining these multiple constraints
enables us to gain a unified understanding of the
magma plumbing system of Asama Volcano from
a crustal magma chamber up to the surface.
A magma chamber horizontally offset from
the summit of a volcano could be a ubiquitous
feature. In fact, tiltmeters in Kirishima Volcano,
SW Japan, tilted towards a point approximately
8 km to the west of the summit in response to an
eruption on 26 January 2011. This is consistent
with the location of a deflation source derived
from GPS observations. Although no seismic velocity models are available to verify whether this
deflation source is characterized by a low-velocity
body, these tilt observations demonstrate the ubiquity of a horizontally offset magma chamber and
structurally controlled magma pathway beneath
active volcanoes.
We thank Japan Meteorological Agency, National
Research Institute for Earth Science and Disaster Prevention of Japan, and Geospatial Information Authority of
Japan for access to their data. Comments by D. Pyle,
C. Wauthier, and an anonymous reviewer improved the
manuscript. This research was supported by the Grantin-Aid for Scientific Research (21740319, 22540431,
22740289) from Japan Society for the Promotion of
Science.
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