Volcanic deformation of Atosanupuri volcanic complex in the Kussharo caldera, Japan, from 1993 to 2016 revealed by JERS-1, ALOS, and ALOS-2 radar interferometry
© The Author(s) 2017
Received: 7 March 2017
Accepted: 29 May 2017
Published: 8 June 2017
Observation of crustal deformation around a volcano provides important information about the dynamic processes of the transfer of volcanic fluids that could lead to an eruption. Continuous GNSS (Global Navigation Satellite System) measurements at permanent sites are now used commonly for monitoring the crustal deformation of volcanoes. However, volcanic deformation is frequently localized within a small area and thus, a very dense network of GNSS receivers would be necessary to capture all the deformation features. It is often difficult to achieve a GNSS network with such density because of the ruggedness of terrain, lack of prerequisite infrastructure such as electricity and cabling for signal transmission, and harsh conditions caused by volcanic activity. L-band synthetic aperture radar (SAR) interferometry (InSAR) has many advantages in relation to volcanic monitoring, and it has become a standard tool for monitoring deformation of the Earth’s surface (e.g., Chaussard et al. 2013; Fournier et al. 2010; Fujiwara et al. 1998). This remote sensing technique is capable of mapping deformation fields over large areas to a high degree of accuracy (i.e., a few centimeters) without need for ground-based measurement facilities. The Japanese Earth Resources Satellite (JERS-1) launched by NASDA [now the Japan Aerospace Exploration Agency (JAXA)] is one such satellite equipped with L-band sensors, and it acquired a large volume of data globally during 1992–1998. To verify the capability of L-band InSAR when applied to volcano monitoring, we performed InSAR analyses using improved processing algorithms (Fujiwara et al. 1998; Tobita et al. 1998) to discover magmatic crustal deformation that could not be detected using other geodetic techniques.
Data and methods
List of interferograms analyzed
Acquisition date master image
Acquisition date slave image
Time span (day)
Maximum displacement (cm)
Corrected maximum displacement (cm)
May 18, 1993
May 9, 1997
August 13, 1993
April 21, 1995
August 13, 1993
July 22, 1998
September 13, 1994
August 31, 1995
September 13, 1994
April 25, 1998
April 21, 1995
April 7, 1996
April 21, 1995
July 22, 1998
July 19, 1995
June 9, 1998
July 19, 1995
July 23, 1998
August 31, 1995
October 31, 1997
August 31, 1995
April 25, 1998
April 7, 1996
March 25, 1997
March 25, 1997
June 8, 1998
June 9, 1998
July 23, 1998
September 22, 2007
August 15, 2010
May 9, 2008
November 15, 2010
August 9, 2014
August 6, 2016
August 25, 2014
June 13, 2016
September 29, 2014
July 4, 2016
For geodetic inversion, we used the formulas reported by Okada (1985) to calculate surface deformations by a source buried in a homogeneous elastic half-space. We followed the nonlinear inversion method of Matsu’ura and Hasegawa (1987) to estimate the parameters of a pressure source. We tested two different pressure sources: (1) a point source (Mogi 1958) and (2) a horizontal planar source (sill). In some cases, results may depend on the input initial value of the model simulation. For the Mogi source, results of depth depend on the initial volume; therefore, we selected the initial volume value to minimize the Akaike information criterion (AIC).
Figure 1 shows a typical deformation field mapped by InSAR from August 1993 to April 1995 along with the topography of the study area. The area of changing colors at the center of the figure depicts the deformation of the ground surface. The horizontal dimension of the deforming area is about 10 km in diameter, and it covers almost the entire area of the Atosanupuri volcanic complex. The maximum range change due to surface deformation was observed near the center of the deforming area, and it reached approximately 20 cm displacement in the direction toward the satellite (upward and/or east-southeastward displacement). In this study, we interpreted that observed radar LOS displacements indicate crustal uplift/subsidence of the Atosanupuri volcanic complex (e.g., Fialko and Simons 2001). As a matter of course, we paid attention to the fact that LOS displacement includes horizontal components (e.g., Segall 2010).
Net deformation in mid-1998 was about half the peak in 1995. It is interesting to note that an earthquake swarm was observed in 1994 (Earthquake Research Institute 2002). Earthquake depths were less than 10 km, and the largest magnitude of this swarm event was 3.2. Since 1950, several earthquakes > M5.0, including an M6.5 event in 1967, have been recorded near the volcano (Katsui et al. 1986). However, the epicenters of these large earthquakes were distributed within an area about 10 km southwest of the volcano. The earthquake swarm in 1994 was unique because the epicenters were just beneath the Atosanupuri volcanic complex (Motoya and Ichiyanagi 1996). Another small earthquake swarm was recorded in 1998 (Fig. 3), and it is interesting that the only uplift observed after 1995 occurred from June to July 1998.
Calculated volcanic source parameters
rms of residuals (cm)
5.5 ± 0.1
2.8 × 107
6.0 ± 0.2
8.4 ± 0.4
1.9 ± 0.7
12.2 ± 2.9
1.4 ± 0.6
2.2 × 107
5.3 ± 0.2
7.1 ± 0.4
1.4 ± 0.7
5.0 ± 2.9
−1.0 ± 0.5
−1.0 × 107
We also inverted the subsiding deformation from July 1995 to June 1996 to estimate all the parameters. We found the horizontal position and the shape of the sill during subsidence to be almost identical to those during uplift. Although a difference (~700 m) in the depths of the sills between subsidence and uplift was found, it was probably attributable to interferogram errors of several centimeters. This finding indicates the position of the source remained static throughout the entire uplift–subsidence episode during 1993–1998.
Uplift and subsidence series
We also recalculated the volcanic source parameters shown in Table 2 using the corrected interferograms. In the calculation of the corrected interferograms, the available data area was reduced because of the averaging of several decorrelated interferograms. Then, we inverted the deformations to estimate the open parameter of the sill. After the correction of the small-scale deformations, the open parameter in Table 2 changed by +0.05 m for the uplift interferogram and by −0.01 m for the subsidence interferogram, and we confirm these values are within the error level. The modeled depth of the source and the increase in volume were similar to those of the Three Sisters volcanic center in the USA. (Wicks et al. 2002).
There are two possible mechanisms for the uplift: (1) migration of magmatic fluids from deeper to shallower parts (Chang et al. 2010; Wicks et al. 2006) and (2) hydrothermal fluid (water and gases) activity (Williams-Jones et al. 2003). However, hydrothermal fluid activities are often found at depths <5 km in other volcanoes such as the Campi Flegrei caldera in Italy (Battaglia et al. 2006) and by numerical simulations (Hutnak et al. 2009). Honda et al. (2011) reported that a low-resistivity (~10 Ωm) structure exists just under the Atosanupuri lava dome and that melts with very low resistivity (~1 Ωm) probably exist at or deeper than 14 km. As the upper depth of the low-resistivity structure is approximately 6 km, the magma probably climbed to the depth of the sill and then migrated laterally, forming a new sill or opening a pre-existing sill. The increase in seismicity can also be explained by changes in the stresses induced by inflation of the sill. We infer that in late 1993 or early 1994, the magma that likely gained buoyancy by some mechanism, presumably via vesiculation, began to ascend from deep below the surface. The total magma volume that contributed to the inflation of the sill was more than 2 × 107 m3 by the middle of 1995 (Table 2). Although the interferograms do not show deformation of the deflating magma reservoir, they do show there was no lateral transport of mass. In the case of a propagating dyke coupled to a magma chamber, Irwan et al. (2006), Rivalta (2010), and Segall et al. (2001) showed that the volume of an inflating dyke is more than three times that of a deflating spherical magma chamber; therefore, the magma probably came from just under the sill and the deformation of the deflating magma reservoir was hidden or overlapping within the deformation field of the inflating sill.
The subsidence began around the summer of 1995. By the summer of 1998, the net volume of the sill during the entire process decreased to about 1 × 107 m3 (Table 2), which is approximately half the volume of the inflation that occurred during 1993–1995. The model simulation (Table 2) suggests that uplift and subsidence have the same source; therefore, there are several possible mechanisms for the subsidence: (1) drainback of the magma; (2) contraction of hot crystalline rock during cooling; and (3) densification of a cooling magma body during crystallization (Dzurisin et al. 2002). The contraction of rock during cooling cannot account fully for the deformation. The coefficients of the thermal expansion for gabbro and granite under standard conditions are 1.6 × 10−5/ °C and 2.4 × 10−5/ °C (Turcotte and Schubert 1982). If the temperature decreased by 1000 °C, a volume change of only several percent would occur. The third mechanism, crystallization, explains 20% of the volume change at most (Caricchi et al. 2014). Viscoelasticity might explain that the deformation shown in Fig. 9 relaxed exponentially with time (Yokoyama 2013) and that the viscoelasticity affected both the uplift and the subsidence. Dragoni and Magnanensi (1989) highlighted that seismic activity began after the onset of ground uplift in the Campi Flegrei caldera because of viscoelasticity; however, the two earthquake swarms in the Kussharo caldera each accompanied or slightly preceded the uplifts. Although the combination of cooling, crystallization of the magma, and viscoelasticity might have contributed to part of the subsidence, a downward drainback of magma remains the most probable mechanism for the subsidence.
As the horizontal distribution of the lava domes of the Atosanupuri volcanic complex and the top of the low-resistivity structure almost match the dimensions of the sill, we speculate the sill likely persists there and that it has been operating as a secondary magma reservoir of the volcanic complex from the past. Similar changes from uplift to subsidence have been observed in the Yellowstone caldera, and Wicks et al. (2006) inferred continuous movement of molten basalt into and out of the Yellowstone volcanic system. As our study lacks concrete evidence supporting the assertion of magma withdrawal as the cause of subsidence, measurements of microgravity change (de Zeeuw-van et al. 2005; Williams-Jones et al. 2003) would support such an analysis in future monitoring of the volcano.
The above scenario raises two questions. (1) Why was the ascent of magma blocked at the depth of 6 km, forcing the magma to migrate laterally? (2) Why did subsidence of the volcano follow the uplift? One possible answer to the first question is that there might be a vertical density gap at the depth of 6 km and that the buoyancy of the magma (Taisne et al. 2011; Wicks et al. 2006) might have been sufficient to rise through the lower and denser mother rock, but not to ascend through the lighter upper layer. On the other hand, a change in fracture orientation and/or stress conditions might have allowed lateral intrusion of the sill instead of vertical migration (Taisne et al. 2011). Thus, as the supply of new magma from deep below the surface continued, there was no other route for the stagnant magma, which was blocked from rising further, other than to migrate laterally. The answer to the second question might be that the initial rate of magma supply from depth decreased and that gradual degassing from the magma remaining in the sill after the ascent reduced its buoyancy, which resulted in the descent or drainback of magma to depths further below the surface (Moran et al. 2011).
The InSAR observations revealed that two types of deformation occurred silently under the Atosanupuri volcanic complex. One type comprised the deformation sequence from uplift to subsidence during 1993–2016. As the deformation had a similar shape throughout the sequence, the deformation could be attributed to a common source. The second type was the small constant subsidence of some lava domes with small-scale patterns throughout the three periods 1993–1998, 2007–2010, and 2014–2016. At the time of deformation, the volcano was not under observation by either the GNSS or any other continuous crustal deformation monitoring system (e.g., tiltmeters or extensometers); therefore, this volcanic event remained unrecognized until the retrospective satellite analysis presented in this study revealed the associated deformation. As this deformation occurred in a volcanic dome that does not show remarkable active volcanic phenomena, continuous monitoring of the Atosanupuri volcanic complex is essential for assessing the potential for future eruptions.
The use of three different SAR satellites meant it was possible to detect the long-lasting crustal deformation sequence. However, the absolute changes in the long-term deformation process could not be established because the interferograms of each of the three satellites were not directly comparable. We hope that future SAR satellites will have capabilities that allow the construction of interferograms comparable with past satellites, and that they will help uncover volcanic activities from the past to the future.
SF analyzed the interferograms and drafted the manuscript. MM and TN constructed the synthetic models, and MT, HY, and TK supported in the construction of the interferograms. All authors read and approved the final manuscript.
The JERS-1 SAR data used in this study were provided by JAXA. Ministry of Economy, Trade and Industry of Japan (METI) and JAXA retain the ownership of the original data. The Geospatial Information Authority of Japan (GSI) obtained ALOS/PALSAR data through the “Joint Cooperative Agreement between GSI and JAXA for observation of geographic information using ALOS data.” Ownership of the ALOS/PALSAR data is retained by JAXA and METI. ALOS-2 data were provided under a cooperative research contract with JAXA. Ownership of ALOS-2 data is retained by JAXA. We used the earthquake data file “Japan University Network Earthquake Catalog Hypocenters File.” Hypocenter data processed by the Japan Meteorological Agency (JMA) were used.
The authors declare that they have no competing interests.
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