Excitation mechanism of large-amplitude volcanic tremor associated with the 1989 Ito-oki submarine eruption, central Japan

Journal of Volcanology and Geothermal Research, 55 ( 1993 ) 33-50

33

Elsevier Science Publishers B.V., Amsterdam

Excitation mechanism of large-amplitude volcanic tremor associated with the 1989 Ito-oki submarine eruption, central Japan Motoo Ukawa National Research Institute for Earth Science and Disaster Prevention, Tennodai 3-1, Tsukuba, Ibaraki 305, Japan (Received November 27, 1991; revised version accepted August 5, 1992 )

ABSTRACT Ukawa, M., 1993. Excitation mechanism of large-amplitude volcanic tremor associated with the 1989 Ito-oki submarine eruption, central Japan. J. Volcanol. Geotherm. Res., 55: 33-50. Unusually large-amplitude volcanic tremor was observed in association with the 1989 Ito-oki submarine eruption off the east coast of Izu Peninsula, central Japan. The largest-amplitude tremor occurred on July 13, 1989, and was studied by using digital seismic data (National Research Institute for Earth Science and Disaster Prevention), in the distance range of 6-250 kin. Spectral analysis shows that the seismograms have two kinds of predominant spectral peaks, low-frequency peaks around 1 Hz and middle-frequency peaks between 2 and 7 Hz. The relative amplitude ratios of these peaks show distinctive spatial variation depending primarily on the distances from the crater. Low-frequency waves are predominant at close stations (distances less than 50 km from the crater) similar to typical low-frequency volcanic tremor, while at more distant stations the middle-frequency waves are more prominent than low-frequency waves. Propagation velocity analysis for narrow-frequency bands on the basis of correlation of envelopes of seismograms reveals that the middlefrequency waves are compressional body waves and the excitation of shear wave was ineffective, suggesting an explosive or implosive source as the simplest source model. In contrast to the middle-frequency waves, our analysis showed no evidence for the low-frequency waves to have propagated with any body wave velocity. Consideration of the spatial pattern of both waves suggests that surface waves were trapped in sedimentary layers from the shallow source and that no lowfrequency oscillations dominated at the source.

Introduction

Unusually large-amplitude continuous tremor was observed during a seismo-volcanic activity occurring off the east coast of Izu Peninsula, central Japan, in July 1989. Some of the tremor events were felt in a nearby city, Ito, causing public awareness. During the largestamplitude tremor, a submarine eruption was observed on the sea surface at 3 km northeast of Ito city (Fig. 1 ). The tremor signals were Correspondence to: M. Ukawa, National Research Institute for Earth Science and Disaster Prevention, Tennodai 3-1, Tsukuba, Ibaraki 305, Japan.

detected at seismic stations at distances more than 300 km from the eruption site (Japan Meteorological Agency, 1990). The purpose of this study is to understand the relationship between the eruption and the large-amplitude tremor. The region where the eruption occurred has experienced seismic swarms in this century. The seismic swarms were first recorded by seismometers in 1930 (Yoshida and Hamada, 1991 ), and, after a long quiescence, resumed and have been periodically active since 1978 (Ishida, 1987). Geologically, the eastern side of Izu Peninsula is known as the region of Higashi-Izu monogenetic volcano group, and

0377-0273/93/$06.00 © 1993 Elsevier Science Publishers B.V. All rights reserved.

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more than 70 monogenetic volcanoes on land and more than 40 in the sea region have been identified (Aramaki and Hamuro, 1977; Hamuro et al., 1980). The seismic swarms had been thought to be one of the crustal activities associated with formation of a new monogenetic volcano (Kuno, 1954; Ishida, 1984). The present seismo-volcanic activity occurred at almost the center of the seismic network deployed by the National Research Institute for Earth Science and Disaster Prevention (NIED) that provided digital seismic data of good quality. As we will show later, seismograms of the tremor exhibit wide variability in the predominant frequency, depending primarily on the distances from the crater found at the eruption site by the Hydrographic Department of Maritime Safety Agency (HDMSA), Japan (Oshima et al., 1991 ). Sta-

tions within ~ 50 km from the crater show a low-frequency signature with a predominant frequency of 0.5 to 1 Hz, typical for low-frequency volcanic tremor, while at more distant stations seismic waves with frequency of ~ 5 Hz are prevailing and the low-frequency signature is less predominant. The variability may be closely related to an excitation mechanism of the tremor. In this paper we attempt to reveal the excitation mechanism of the tremor by distinguishing between propagation effects and source excitation. Several studies have been performed regarding the tremor activity associated with this eruption: Kudo et al. (1991) analyzed seismograms of strong-motion accelerographs recorded at close-distance stations and found that the tremor was characterized by high frequency of water waves and low frequency of surface waves. Goto et al. ( 1991 ) found that the tremor has three spectral peaks, around 1 Hz, 7 Hz and 20 Hz, on the basis of temporal seismic stations of JMA. Yamasato et al. ( 1991 ) concluded on the basis of seismograms of JMA seismic stations at close distances that the tremor is composed of high-frequency (more than 3 Hz) P waves and low-frequency surface waves, both of which were associated with explosions, and continuous low-frequency tremor composed of Love waves. In this study we use seismograms of more than 50 stations widely distributed in the Kanto-Tokai area in the distance range of 6 to 250 km.

Chronology of the 1989 Ito-oki submarine eruption The 1989 Ito-oki submarine eruption involved various crustal activities including seismic swarms, low-frequency earthquakes, continuous tremor events and crustal deformation. We briefly summarize the chronology of the activities to set the stage for the occurrence of the strong tremor in conjunction with the seismo-volcanic activity. Crustal activities associated with the present

35

E X C I T A T I O N M E C H A N I S M O F L A R G E - A M P L I T U D E V O L C A N I C T R E M O R , 1989 ITO-OKI E R U P T I O N

tions, stating that a major magma intrusion into the shallow part of the crust took place from July 4 to July 10. The length and opening of the dyke are estimated at approximately 4 km and 1 m, respectively. After cessation of swarm activity on July 10, low-frequency earthquakes occurred occasionally (Yamasato et al., 1991 ), as indicated in Figure 2. Since July 11, continuous tremor was observed intermittently. In Figure 2, activity of the continuous tremor is indicated with their duration time measured on seismograms recorded at HTS, the closest station among the NEID seismic network, located 6.6 km northeast from the eruption site (Fig. 1 ). The longest tremor occurred on July 12, lasting more than 1 hour and the largest-amplitude tremor took place from 18:33 to 19:20 (Japan Standard Time, JST) on July 13. On that day, water domes and water plumes were observed on the sea surface from the nearby coast and more closely from a survey vessel of HDMSA (Oshima et al., 1991 ). This observation confirmed that the sequence of crustal activity culminated in a submarine eruption.

eruption are summarized in Figure 2, which demonstrates daily earthquake numbers, lowfrequency earthquake activities, continuous tremor events and a period of large crustal deformation. The seismic activity began on June 30, 1989. The number of earthquakes greatly increased on July 4 and decreased with several intermittent peaks up to July 10, including the largest event, an M 5.5 earthquake on July 9. Seismograms of the swarm show clear P and S phases, indicating that these swarms are of tectonic type. For most of the swarms, the focal mechanism was of double couple type, the strike-slip fault with P axis in a NW-SE direction (Matsumura et al., 1991 ). During the period of intensive seismic swarm from July 4 to July 10, large crustal movements were observed by various measurements; a tiltmeter (Yamamoto et al., 1991 ), GPS (Shimada et al., 1990 ), triangle measurements (Tada and Hashimoto, 1991) and a laser distance measurement (Tsuneishi, 1991 ), indicating a NE-SW extension. Okada and Yamamoto (1991) proposed a model of magma intrusion to explain these observa-

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HDMSA conducted sea bottom topography surveys just before and after the eruption of July 13, and detected a dome-shaped bulge, 25 m high, five minutes before the eruption, and a new crater with a diameter of approximately 200 m, on July 15 (Oshima et al., 1991 ). It should be emphasized that no significant crustal deformation was detected in association with the tremor activity and the eruption observed on the sea surface (Okada and Yamamoto, 1991 ), suggesting that the eruption and the tremor events were relatively small and shallow in contrast to the magma intrusion generating the seismic swarm activity. Data In the following analyses, we used seismic data obtained by the NIED seismic network. The station distribution is shown in Figure 1. The NIED seismic network has 76 stations, with an average spacing of approximately 20 km, covering central Japan (Fig. 1). Threecomponent velocity-type seismometers with a natural frequency of I Hz and a damping factor of 0.7 are installed at most of the stations. The seismic data are digitally telemetered to the headquarters of the NIED in Tsukuba. The total dynamic range is 78 dB and the sampling rate is 80 Hz. The three-component digital data are recorded on optical disks using a triggering algorithm. Seismograms of vertical component are continuously recorded on pen-recorders for monitoring purpose. As this seismic network is designed for the study of earthquake prediction, the triggering algorithm is not suited for tremor recording. It, therefore, recognizes tremor as noise. In the present case, unfortunately, the tremor events, except for the largest-amplitude one, were not recorded digitally. Our analysis, therefore, is limited solely to the largest- amplitude tremor, during which the submarine eruption was observed as indicated in Figure 2. For this tremor, digital seismograms of at least the vertical component are available at 58 stations with 1-

Hz seismometers in the distance range from 6 to 250 km. Figure 3 shows a part of continuous monitoring chart recorded at FJM, a relatively low-amplification station located 55.7 km NW from the crater (Fig. 1 ). The tremor began at 18:33 on July 13 and lasted for about 50 minutes, sustaining high amplitude for about l 5 minutes from the onset. Although the amplitude of the tremor was fairly small from 18:50 to 19:03 at FJM, closer stations continuously recorded tremor signals during this period. The tremor signal was distinguished on the seismograms until 19:20. We will focus our analyses on the period from 18:38 to 18:41, because the tremor signals in this period were large enough to trigger the recording system at stations that are more than 200 km from the crater and not too large to clip seismograms at close stations. Three-component (vertical, radial and transverse) seismograms at 20 stations are plotted in Figure 4 for 45 s from 18:38:00, aligned in order of distance from the crater. The latter two components are synthesized from the two horizontal components (N-S and E-W). The stations are selected by the following criteria: spatial distribution is as even as possible, three-component l-Hz seismograms are available, and seismograms are not clipped. Although vertical and E-W component seismograms of HTS are clipped and its N-S component was in trouble of seismometer, we plotted the seismograms in Figure 4 because HTS is the nearest station among the NIED seismic network. We substituted the original N-S and E-W components for the plots of the transverse and radial components, respectively, because we could not correctly transform the seismograms for HTS. The most distinctive feature of Figure 4 is a variability of waveforms, depending primarily on distances from the crater. On the vertical component seismograms (Fig. 4a), predominant frequencies of the tremor at stations with distances less than about 50 km are generally low, ranging 0.5 to 1.0 Hz. The seismograms of

EXCITATIONMECHANISMOF LARGE-AMPLITUDEVOLCANICTREMOR, 1989ITO-OK1ERUPTION 18:30

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close stations, for example NRY, HDA and SMD, exhibit a sinusoidal low-frequency signature just like typical low-frequency volcanic tremor seismograms that appear on many volcanoes. On seismograms of the stations with distances larger than about 50 km, low-frequency waves are less predominant and higherfrequency waves with frequencies of about 5 Hz superpose on the low-frequency waves. On some stations (e.g., YMK, AKW and H H R ) in the distance range 50 to 100 km, higher-frequency waves (2-7 Hz) are predominant over low-frequency waves. In this paper, we refer to frequencies of 0.5-Hz as the low-frequency range and to frequencies of 2-7 Hz as middle frequency for convenience. On the horizontal component seismograms, variability in predominant frequency is less distinctive than on the vertical component at all distances (Fig. 4b and c ), where most of the stations exhibit a low-frequency signature rather than middle-frequency waves.

Observed spectra Spectra of the tremor were estimated for vertical, radial and transverse component seismograms by using the fast Fourier transform method. As can be seen in Figure 4, a predominant frequency of each seismogram fluctuates temporally, even over the short time scale of 45 s. We, therefore, evaluate the spectra by applying a moving time window of 6-s width, making it possible to evaluate spectra in the frequency range of 0.2 to 15 Hz. We move the time window from 18:38:00 to 18:40:00 with no overlapping. The spectra are not corrected for instrumental response. The spectra of observed seismograms were evaluated for 58 stations, of which three-component spectra are available for 44 stations (Fig. l ). Examples of the spectra are plotted in Figure 5, demonstrating the typical spectral characteristics depending primarily on distance from the crater as mentioned in the pre-

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ones are much smaller than the observed ratios, maybe less than 10 %. Spectra of the stations in the distance range 50 to 100 km, for example YMK, AKW and H H R in Figure 5, are quite different from those of closer stations. Several peaks with comparable amplitudes exist between 2 and 7 Hz, and the strongest peak appears between 3 and 4 Hz for the vertical component at AKW and HHR. On plots of the averaged spectra, the peak amplitudes in the middle-frequency range are larger than (or comparable to) those at low frequency. In contrast to the spectra of vertical components, the spectral peak around 1 Hz is prominent for the horizontal components, but not so much as for the closer stations. At stations with distances more than 100 km from the crater, we recognize several spectral peaks in the middle-frequency range. Although their amplitudes are generally less than the peaks at

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low frequency, large variability among stations is seen in amplitude ratios of low and middlefrequency waves. The most important feature in Figure 5 is that frequencies of the most prominent spectral peak show spatial variability, depending primarily on distance from the crater. The ten-

dency for the stations closer to the crater to exhibit a prominent peak at lower frequency as opposed to distant stations suggests that the spatial variability of predominant frequency cannot be explained by wave attenuation during propagation. Comparing three-component seismograms of the same station, amplitudes

of spectral peaks at the middle-frequency range relative to those at low frequency are smaller for the horizontal components than for the vertical components in all distance ranges.

Propagation velocities Propagation velocity is, in general, a useful indicator to identify wave types. For continuous tremor, however, it is difficult to estimate propagation velocities, partly because of the difficulty in identifying the onset time for each phase and partly because of superposition of various types of seismic waves. For continuous seismic waves, like surface waves and microseisms, a beam-forming method (Capon et al., 1969 ) can be applied to estimate phase velocities using a dense seismic array. In applying this method to seismograms recorded by the NIED seismic network, we need to make some modifications, because the average spacing of the stations (20 km ) is far larger than the typical wavelength of either body waves or surface waves in frequency ranges higher than l Hz. We evaluate propagation velocities on the basis of correlation of envelopes of the tremor seismograms in place of original seismograms themselves, using the concept of the beamforming method explained below. This method is suitable for our data because the amplitude of the envelope presumably has a longer correlation length than the waveform itself, indicating propagation of wave packets. For seismic waves excited continuously with small fluctuation of amplitude, this method is ineffective, because peaks of envelop seismograms may not be large enough to obtain high signalto-noise ratios. In practice, the tremor analyzed in this study shows large fluctuations of amplitude in both ranges of the low and middle frequencies, indicating that this method is effective for the present data set. The tremor signal of the i-th station is expressed by xi (t). The seismogram which has been processed through band-pass filters (cutoff frequencies,f~ andf2,f~
by xi (t;f~,f2). Each filtered seismogram is normalized by its maximum amplitude. We obtain a seismogram envelope by squaring the seismogram, corresponding to an energy seismogram, and then smoothing by a moving average box-car window with 0..5-s width. The envelope seismogram is expressed as y~(t;./i./~ ) as shown in equation ( 1 ):

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where Tw is the window width, dt is a sampling interval, and Z is a summation over discrete time t within the window. A beam seismogram, S(t; v, f~, f2), is constructed by summing up y~(t;f~,f2) over all stations with an assumed propagation velocity, v:

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where & is a distance from the crater to the i-th station, and N is a total number of stations. To search a velocity which maximizes a peak amplitude of each beam seismogram, we calculate:

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where Say is an averaged amplitude of S(t; v, f~, f2) over the whole period. When v differs largely from a true propagation velocity, beam seismograms do not have any strong peak, showing many small peaks and troughs, and making a fluctuation from the average, A (t; v, fl,f2 ), small. When v is close to a true velocity, wave packets with high amplitude result from constructive interference in relatively small time spans, whereas destructive interferences reduces amplitudes of S( t; v, f~, f2 ) outside of the wave packets, thus making a maximum of A ( t; v, f~, ~ ) large. We can, therefore, use A (t; v, f~, ~ ) as a useful indicator for wave packet construction. The maximum value of A ( t; v, f~, f2 ), over the time span for which S is calculated, is referred to as Amax. Amaxis a function of v if the cut-off frequencies are fixed. We define a propagation velocity as the velocity which gives the m a x i m u m Amax. If body waves dominate, the

EXCITATION MECHANISM OF LARGE-AMPLITUDEVOLCANICTREMOR, 1989ITO-OKI ERUPTION

estimated values of the propagation velocity may be constant, showing crustal seismic velocities. If surface waves with dispersion are predominant, the propagation velocity estimated by this method should be a function of f~ and f2. In practice we chose 8 frequency ranges, a low-pass filter with 1-Hz cut-off and 7 bandpass filters with 2-Hz band widths from 1 to 15 Hz. Amax was searched for velocities from 3 to 9 k m / s with 0.25 k m / s increment, this velocity range covering typical crustal P and S wave velocities. Because of the limited length of seismograms recorded continuously, we cannot expand the analysis below 3 km/s. The above velocity range is not enough to cover surface wave velocities with frequencies of around 1 nz, whose propagation velocities are expected to be lower than the S wave velocity in the very shallow part of the crust, 2-3 k m /

43

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We applied this method to the vertical and transverse component seismograms for two periods (30 s long and 60 s long) from 18:39:10 at A = 0 km, using stations with 1-Hz seismometers (58 vertical components and 44 transverse components). We mainly show the resuits of the 30-s period and use those of the 60s period to examine stability of the estimated velocities. Examples of the seismograms are seen in Figure 4 a and c. Figure 6 shows examples of A (t; v, fl, f2 ) obtained for vertical seismograms with f~ = 3 Hz and f2 = 5 Hz, that is the frequency range in which the strong peaks of the middle-frequency waves appear. Figure 6 indicates two clear and one small wave packets on the beam seismograms for propagation velocity around 6.0 km/s.The two large peaks correspond to wave groups with the apparent velocity of ~ 6 k m / s appearing on seismograms in Figure 4a. The optimum propagation velocity was found to be 5.50 k m / s after a full calculation in this case. Figure 7 summarizes the resulting Am,x versus propagation velocity for all frequencies in the range for vertical seismograms. In this fig-

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ure, triangles indicate velocities giving the maximum Amax, which are defined as the propagation velocity. In the frequency range higher than 1 Hz, each curve has one peak in the velocity range from 5.25 to 6.5 km/s for both analysis periods, the velocity range of which is consistent with the typical upper crustal P wave velocity, 5.9 km/s, in this region (Asano et al., 1985). The peaks are clear especially for the frequencies between 1 and 7 Hz, which correspond to the middle-frequency range. In contrast to the results for the frequencies higher than 1 Hz, no clear peak was identified for the one with the low-pass filter of 1 Hz, suggesting that low-frequency wave energy propagating with velocities from 3 to 9 km/s is too small to identify and that the low-frequency waves with large amplitudes are incoherent. The small differences between the propagation velocities es-

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timated for 30-s and for 60-s periods indicate the stability of the estimation, ranging within 0.75 km/s. To examine correlation of waveforms on original filtered seismograms, Figure 8 compares the low-pass (1 Hz) filtered seismograms with the band-pass (3-5 Hz) filtered ones. Clear wave packets can be seen on the latter, while it is difficult to distinguish any correlation in the former. From the results obtained above we conclude that wave packets are composed of P waves with frequencies higher than 1 Hz, and that seismic waves corresponding to the spectral peak in the low-frequency range are not P waves travelling through the crust. Figure 9 shows the similar results for transverse seismograms to Figure 7 for vertical seismograms. The peaks of Am~-velocity curves in

< 5s >

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Fig. 8. Original filtered seismograms. (a) 1-Hz low-pass filtered seismograms reduced by 7.5 k m / s , and (b) 3-5-Hz bandpass filtered seismograms reduced by 5.5 k m / s .

EXCITATION MECHANISM OF LARGE-AMPLITUDE

VOLCANIC TREMOR,

Transverse component

011

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45

1989 ITO-OKI E R U P T I O N

Discussion

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Excitation of low-frequency waves

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Velocity, km/s Fig. 9. Amax versus assumed velocity for all frequency ranges for transverse component seismograms.

The excitation mechanism of the low-frequency waves is an important question. Although the low-frequency waves around 1 Hz are generally prominent in the nearer stations, substantial spatial variability exists. To quantify the spatial variability we introduce a ratio R =Am/A1, where Al and A m a r e m a x i m u m amplitudes of seismograms filtered through frequency bands corresponding to the peaks of the low-frequency range and of the middle-frequency range. We chose a frequency band of 0.5-2.0 Hz for the low-frequency spectral peaks and that of 3.0-6.0 Hz for the middle range. Peak amplitudes were picked up on seismograms for a one-minute-period from 18:38:00. The resulting amplitude ratios range from 0.1 to 1.7 as plotted in Figure 10. The plots show that small values of R, indicating relatively large amplitude of the low-frequency waves, 37°

Figure 9 are less prominent than those in the case of vertical component for both periods of 30 s and 60 s. Furthermore, the estimated propagation velocities range between 5.25 and 6.75 k m / s for most of the frequency ranges, indicating that there is no clear evidence of shear wave propagation, even on the transverse component seismograms. Although for the bands of 0-1 Hz, 5-7 Hz and 7-9 Hz, some peaks appear between 3.25 k m / s and 3.75 k m / s, which is consistent with the crustal S velocity, there is no case where the propagation velocities for both periods indicate crustal S velocity. Therefore, it seems difficult to conclude that the peaks showing the crustal velocity represent the source radiation of S waves. We infer that the wave packets are composed of mainly compressional waves and that the seismic waves around 1 Hz did not propagate either with P wave or S wave velocities typical of the upper crust.

35~

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33° N 137° E

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139°

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Fig. 10. Spatial distribution of ratios R=Am/A~(see text). Areas with altitude higher than i 000 m are indicated by hatch.

46

appear in the coastal regions and large ones in mountainous regions. A possible cause of this spatial variability is a combination of an excitation of surface waves at the source and an effective trapping of surface waves in shallow sedimentary layers. Several observations suggest a shallow depth for the tremor source which would support effective excitation of surface waves: T phases were observed at stations near the coast during the tremor period ( Kudo et al., 1991; Ukawa et al., 1991 ); some events occurring during the tremor were located at very shallow depths (0 to 2 km ) ( Goto et al., 1991 ); the tremor accompanied the surface volcanic event of submarine eruption, but not occurred in the swarm period of July 4 to 10 when the large crustal deformation took place (Okada and Yamamoto, 1991 ). For the seismic sources in sedimentary layers with low velocity, strong excitation of surface waves can be expected. Strong motion records of the tremor near the crater show successive occurrences of many impulsive events whose average intervals are approximately 1 to 2 s (fig. 3 in Kudo et al., 1991 ). The seismograms at HTS also show the successive middle-frequency events as shown in Figure 4c. Usually, shallow events have long tails, or coda, comprising mainly surface waves.When such events occur successively with intervals shorter than coda duration, the superposed coda waves make the amplitude large, while body waves are isolated due to their short duration times. In the present case, the amplification mechanism from superposition of low-frequency coda was reinforced by excitation of surface waves from the shallow source. In the Kanto-Tokai region, thick sedimentary layers with P wave velocities of 2-3 k m / s exist in the coastal regions (e.g., Asano et al., 1982 ), where some of the NIED seismometers are deployed. On the other hand, in the mountainous region seismometers are positioned on exposed basement rocks with P wave velocities of 5-6 k m / s (e.g., Aoki et al., 1972). The

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Fig. 11. Schematic diagram showing the difference in propagation of tremor signals between the stations in mountainous and coastal regions.

low-frequency waves excited around the source were trapped in the shallow sedimentary layer along the coast, reaching the coastal stations, while in the mountainous region, the low-frequency signals were attenuated, because there was no sedimentary layer providing a channel for surface waves. Figure 11 schematically indicates the difference in propagation characteristics between coastal and mountainous regions. Our conclusion, that the prominent lowfrequency waves are of surface wave origin, is consistent with the results obtained by Yamasate et al. (1991 ) and Kudo et al. (1991) on the basis of polarization analyses of seismograms at close distances, that is, the low-frequency tremor is composed of both Rayleigh and Love waves. Absolute a m p l i t u d e s

In order to calculate amplitude decay of the tremor with distance, we applied a relation between velocity amplitude ( A t ) and distance ( r ) , A t = a r - b where a and b are coefficients to be estimated, to m a x i m u m amplitudes of the vertical seismograms through low-pass filter ( 1 Hz) and those through band-pass filter (3-5 Hz) for the two-minute period from 18:38:00. Figure 12 shows results for both of the fre-

EXCITATION MECHANISM OF LARGE-AMPLITUDE VOLCANIC TREMOR, 1989 ITO-OKI ERUPTION

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Fig. 12. Plots of maximum amplitudes (At) of filtered seismograms versus distance (r) for the period 18:38:00 - 18:40:00. The straight lines show the least-squares fitting of the relation, A t = a r -b. (a) Low-pass filter ( 1 Hz), and (b) Band-pass filter (3-5 Hz).

quency ranges. The coefficients are calculated by applying the least-squares fitting to 58 amplitude data. The exponents, b, calculated for 0-1 H z a n d 3-5 H z a r e 1.5+0.1 and 1.3___0.1, respectively, both of which are larger than the value of 1, which is expected for body waves in a homogeneous medium without attenuation. Since seismic waves of frequency range of 3-5 Hz are mainly composed of P waves, the reason of the exponent larger than 1 is qualitatively explained by attenuation and geometrical spreading effects.

47

Although the present result indicates that the seismic waves of 0-1 Hz are mainly surface waves, whose geometrical spreading factor is generally expected to be less than those of body waves, the resulting exponent is larger than that for 3-5 Hz. This discrepancy may result from large absorption and strong scattering of seismic energy in the very shallow part of the crust, through which surface waves with frequencies around 1 Hz propagated. For volcanic tremor observed in Mount St. Helens, comparable values, 1.5 to 1.7, are obtained for the exponent of the amplitude distance relation (Hofstetter and Malone, 1986). The coefficient, a, corresponds to an amplitude at unit distance, 1 km in the present study. We evaluate reduced displacement (Aki and Koyanagi, 1981 ), which is a useful measure to compare amplitudes of tremor, from the calculated values of a, 1.3× 10 -3 m/s for 0-1 Hz and 2.9X 10 -4 m/s for 3-5 Hz. The reduced displacement is approximately calculated to be ar/ 2x/]rrf, where f is a representative frequency. The reduced displacements are obtained to be 1.5× 103 c m 2 for 0-1 Hz (assuming f = 1 Hz) and 8.2X 10cm 2 for 3-5 Hz ( f = 4 Hz). As the amplitude of the tremor in this period is approximately half of that at the climax (18:42) (Yamasato et al., 1991), the maximum of the reduced displacement may be twice as large as the above values. The reduced displacements estimated for the 1989 Ito-oki eruption are much larger than those obtained for deep volcanic tremor under Kilauea, 3 to 64 cm 2 (Aki and Koyanagi, 1981 ), and those for volcanic tremor of Mount St. Helens in August and October 1980, 1 to 20 cm 2 (Fehler, 1983). The reduced displacement for 0-1 Hz for the 1989 Ito-oki eruption is larger by several times than that of the largeamplitude volcanic tremor of May 18 eruption of Mount St. Helens, 260 cm 2 (Fehler, 1983). Even though different excitation mechanisms may work for these events, we can recognize the large excitation of seismic waves at the

1989 Ito-oki eruption by comparing the reduced displacements.

Source model Unfortunately, we do not have enough information with regard to the source. The most valuable information is the lack of shear wave energy in all directions. It is difficult to explain the lack of shear waves by absorption in magma body because of the total azimuthal distribution of this phenomenon. The simplest source satisfying the ineffective excitation of shear waves is an explosive or an implosive type. The effective excitation of compressional waves and the ineffective excitation of shear waves are consistent with the observation of larger contamination of middle-frequency waves on vertical component seismograms than on horizontal components as shown in Figure 5. Ukawa et al. (1991) suggested that four large-amplitude impulsive events occurring during the tremor at 19:03 to 19:06 can be explained primarily by an implosive source. Since these impulsive events show the similar characteristics of seismograms, the spectra and the ineffective shear wave excitation, to the tremor associated with the eruption, Hashimoto et al. ( 1991 ) and Goto et al. ( 1991 ) suggest that intermittent occurrence of impulsive events similar to those at 19:03 to 19:06 resulted in the large-amplitude tremor during the eruption. It is not certain, however, from the present data whether all the events forming the tremor were excited from an implosive source like the four impulsive events or not, because polarities of the first motions of P waves cannot be examined. In a geological perspective, Ono et al. ( 1990 ) suggested that the explosion on July 13 was magma phreatic-type. The successive occurrences of small seismic events are certainly closely related to the explosion at the sea bottom as suggested by crustal m o v e m e n t measurements (Okada and Yamamoto, 1991). Yamasato et al. (1991) suggested multiple

sources for the tremor, explosions and the other source which continuously excited love waves. Although the spectra observed by the NIED seismic network suggested large excitation of seismic waves with frequencies 2 to 7 Hz, spectra of closer stations show excitation at frequencies of more than 10 Hz (Kudo et al., 1991; Goto et al., 1991 ). In this study we are not able to make a model for the source process because of the paucity of information. It, however, should be emphasized that the seismic energy was mainly radiated in the middle-frequency range of 2 to 7 Hz, which may be related to the source dimension. For the large-amplitude tremor events observed at21 h o n J u l y 11 a n d 9 h t o 1 0 h o n July 12, no digital data were recorded, and only monitoring charts are available. They show spatial variability similar to that we analyzed in this study, suggesting that the same amplification mechanism as for the tremor during the eruption can explain the excitation of lowfrequency waves on July 11 and 12. Another large-amplitude volcanic tremor was reported for the eruption of Shiveluch volcano in 1964 (Gorshkov and Dubik, 1970) and still another for the May 18 eruption of Mount St. Helens in 1980 (Malone et al., 1981 ). Tremor signals reached 430 km in the former case and more than 650 km in the latter. The 1989 Ito-oki submarine eruption excited volcanic tremor comparable to these events in amplitude, but the magnitude of the eruption of the 1989 Ito-oki eruption is far smaller than the others. The volume of deposits is estimated to be 1.5 km 3 for Shiveluch volcano (Gorshkov and Dubik, 1970) and total volume of avalanche of Mount St. Helens is 2.8 km 3 (Christiensen and Peterson, 1981 ), while the 1989 Ito-oki eruption constructed only a small crater with a diameter of 200 m. The quite effective excitation of volcanic tremor at the 1989 Ito-oki eruption may be related to efficiency of seismic wave generation near the sea bottom. Low-frequency volcanic tremor has been

EXCITATIONMECHANISMOF LARGE-AMPLITUDEVOLCANICTREMOR,1989ITO-OKIERUPTION

successfully synthesized from seismic waves emanating from a source predominantly oscillating with low frequency (e.g., Crosson and Bame, 1985; Chouet, 1987). However for a shallow focus tremor, which excites surface waves, we need to examine carefully the possibility that low frequencies result from superposition of surface waves emanating from successive small events. McNutt (1986) has also indicated the importance of surface waves as the source of low-frequency volcanic tremor. The present result suggests the importance of high-frequency observation to locate the source of volcanic tremor, even when low-frequency waves are predominant.

Conclusions The unusually large-amplitude tremor associated with the 1989 Ito-oki eruption was studied by using seismograms recorded by the NIED seismic network with a broad distance range of 6-250 km. The observed spectra show two kinds of peaks, low-frequency peaks around 1 Hz and middle-frequency peaks between 2 and 7 Hz. The ratio of the peak amplitudes exhibits spatial variability. At close stations (within 50-km distance) low-frequency waves dominate, typical for low-frequency volcanic tremor, while many seismograms of stations farther than 50 km from the crater show predominant middle-frequency waves. Propagation velocity analysis for narrow-frequency bands reveals that the middle-frequency waves are composed of compressional body waves and that shear wave excitation was ineffective, suggesting an explosive or implosive source for the middle-frequency waves. In contrast to the middle-frequency waves, the analysis did not show any body wave velocity for the low-frequency waves. Considering the propagation velocities and the spatial distribution of high amplitude of the low-frequency waves relative to the middle-frequency waves, we infer that the low-frequency waves were surface waves trapped in sedimentary layers

49

from a shallow source and that the low-frequency oscillation did not dominate at the source.

Acknowledgement The author would like to thank J.M. Lees for critically reading the manuscript. Comments and suggestions by the anonymous reviewer, K. Horai, H. Yamasato, Y. Okada and S. Sekiguchi were useful to improve the manuscript.

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50 Gorshkov, G.S. and Dubik, Y.M., 1970. Gigantic directed blast at Shiveluch volcano (Kamchatka). Bull. Volcanol., 31: 262-288. Goto, K., Mori, S. and Fukui, K., 1991. Volcanic tremor accompanying the 1989 submarine eruption off east coast of Izu Peninsula. J. Phys. Earth, 39: 47-64. Hamuro, K., Aramaki, S., Kagami, H. and Fujioka, K., 1980. The Higashi-Izu-oki submarine volcanoes, Part 1. Bull. Earthq. Res. Inst., Univ. Tokyo, 55:259-297 (in Japanese with English abstract). Hashimoto, T, Maeda, K., Odaka, T. and Seino, M., 1991. Waveform analysis of volcanic tremors associated with the July 1989 volcanic activity offthe east coast oflzu Peninsula. J. Phys. Earth, 39: 65-78. Hofstetter, A. and Malone, S.D., 1986. Observations of volcanic tremor at Mount St. Helens in April and May 1980. Bull. Seismol. Soc. Am., 76: 923-938. Ishida, M., 1984. Spatial-temporal variation ofseismicity and spectrum of the 1980 earthquake swarm near the Izu Peninsula, Japan. Bull. Seismol. Soc. Am., 74:199221. Ishida, M., 1987. Recent seismic activity in and around the Izu Peninsula. Proc. Symp. Earthq. Predict. ( 1987 ): 51-60 (in Japanese with English abstract). Japan Meteorological Agency, 1990. Seismic activity in and around the Izu Peninsula. Rep. Coord. Comm. Earthq. Predict., 43:140-156 (in Japanese). Kudo, K., Sawada, M., Sakaue, M., Miyazaki, T. and Oshima, S., 1991. Volcanic tremor associated with the 1989 Submarine eruption offlto, Japan. J. Phys. Earth, 39: 27-46. Kuno, H., 1954. Geology and petrology of Omuro-yama volcano group, North lzu. J. Fac. Sci. Univ. Tokyo, Sect. 2, 9: 241-265. Malone, S.D., Endo, E.T., Weaver, C.S. and Ramey, J.W., 1981. Seismic monitoring for eruption prediction. U.S. Geol. Surv., Prof. Pap., 1250: 803-813. Matsumura, S., Obkubo, T. and Imoto, M., 1991. Seismic swarm activity in and around the Izu Peninsula preceding the volcanic eruption of July 13, 1989. J. Phys. Earth, 39: 79-92.

M I K.\vv~: McNutt, S.R., 1986. Observations and analysis of b-type earthquakes, explosions, and volcanic tremor at t'avlof volcano, Alaska. Bull. Seismol. Soc. Am., 76: 153175. Okada, Y. and Yamamoto, E., 1991. Dyke intrusion model for the 1989 seismovolcanic activity offlto, central Japan. J. Geophys. Res., 96: 10,361-10,376. Ono, K., Soya, T., Suto, S., Uto, K. and Yamamoto, T., 1990. Products of the 1989 submarine eruption offlto, east coast of the Izu Peninsula. J. Geogr., 99:142-146. Oshima, S., Tsuchide, M., Kato, S., Okubo, S, Watanabe, K., Kudo, K. and Ossaka, J., 1991. Birth of a submarine volcano "'Teisi Knoll". J. Phys. Earth, 39: 1-20. Shimada, S., Fujinawa, Y., Sekiguchi, S., Ohmi, S., Eguchi, T. and Okada, T., 1990. Detection of a volcanic fracture opening in Japan using global positioning system measurements. Nature, 343:631-633. Tada, T. and Hashimoto, M., 1991. Anomalous crustal deformation in the northeastern lzu Peninsula and its tectonic significance -tension crack model. J. Phys. Earth, 39: 197-218, Tsuneishi, Y., 1991. Continuous monitoring of crustal activity in east lzu Peninsula by automatic electronic distance measurement. J. Phys. Earth, 39:131-140. Ukawa, M., Obara, K. and Fukuyama, E., 1991. Seismic pulses suggesting an implosive source at the 1989 Itooki submarine eruption, central Japan. Geophys. Res. Lett., 18: 873-876. Yamamoto, E., Okada, Y. and Ohkubo, T., 1991. Ground tilt changes preceding the 1989 submarine eruption off lto, Izu Peninsula. J. Phys. Earth, 39: 165-176. Yamasato, H., Yokota, T. and Kashiwabara, S., 1991. Earthquake swarm and volcanic tremors off eastern Izu Peninsula in 1989 -spectral investigation and characteristics of waveforms. J. Phys. Earth, 39: 79-92. Yoshida, A. and Hamada, N., 1991. Redetermination of hypocenters of foreshocks, main shocks, and aftershocks of the Kita-lzu earthquake and the Ito earthquake swarm of 1930. J. Phys. Earth, 39: 329-344.