Seismic attenuation at Rabaul volcano, Papua New Guinea

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Journal of Volcanology and Geothermal Research 130 (2004) 77^92 www.elsevier.com/locate/jvolgeores

Seismic attenuation at Rabaul volcano, Papua New Guinea Ł . Gudmundsson a;1;
, D.M. Finlayson b;2 , I. Itikarai c , Y. Nishimura d , O W.R. Johnson e a

Danish Lithosphere Centre, Copenhagen, Denmark 6 Neilson Street, Garran, ACT 2605, Australia c Rabaul Volcano Observatory, Rabaul, Papua New Guinea d Hokkaido University, Sapporo, Japan Geoscience Australia (formerly Australian Geological Survey Organisation), Canberra, ACT, Australia b

e

Received 16 July 2002; accepted 24 July 2003

Abstract The attenuation structure around Rabaul volcano, New Britain, Papua New Guinea, is studied using broadband records from a pair of sites, inside and outside of the Rabaul caldera complex, using regional earthquakes. Estimates of attenuation for P and S waves between 0.5 and 5 Hz indicate that near-surface rocks within the caldera complex are significantly more attenuative than outside the caldera. The average strength of this anomaly is defined in terms of t*S V0.2 s over and above the region outside the caldera (t*S is the time integral of the inverse quality factor of attenuation, in this case along the path of a shear wave). This attenuation anomaly appears to be strongest at depth and to the south of the Kaivuna site in the northern part of the caldera complex. The ratio of attenuation estimates for P and S waves is Nt*S /Nt*P W2.7, indicating that intrinsic attenuation contributes significantly to the process of attenuation. No frequency dependence of the quality factor is resolved within this frequency band. Path-averaged lithospheric attenuation is estimated from S/P spectral ratios at both stations, yielding quality factors on average QL V180, but possibly significantly lower around and below the volcano. Shear waves are observed along paths passing underneath the central southern caldera at depths greater than about 5 km, although they are strongly attenuated. This indicates that they have not passed through a large, extensively molten body. > 2003 Elsevier B.V. All rights reserved. Keywords: attenuation; caldera; regional earthquakes; magma reservoir

1. Introduction

1

Formerly at the Research School of Earth Sciences, Australian National University. 2 Formerly at Geoscience Australia (Australian Geological Survey Organisation). * Correspoding author. Tel.: +45-3814-2651; Fax: +45-3311-0878. E-mail address: [email protected] (O. Gudmundsson).

The attenuation of elastic waves depends strongly on temperature. Attenuation is quanti¢ed by the quality factor, Q, or internal friction, Q31 , which typically varies by orders of magnitude between ambient temperature and a rock’s solidus (e.g. Kampfmann and Berckhemer, 1985). For shear waves the quality factor must vanish before

0377-0273 / 03 / $ ^ see front matter > 2003 Elsevier B.V. All rights reserved. doi:10.1016/S0377-0273(03)00282-8

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the rock’s liquidus is reached. Seismic attenuation is therefore a useful parameter to characterize the physical state of rocks within a volcano, where melt £uids are thought to reside at shallow levels in the Earth’s crust. A number of factors other than temperature are known to a¡ect the attenuation of seismic waves (e.g. Winkler and Nur, 1982). Of these, rock fractures and hydrothermal £uids are perhaps most important in volcanic settings. Highly fractured rock is more compliant than crystalline rock. Rock that is saturated with a dense £uid would be expected to cause a higher degree of attenuation than if it were ‘dry’. Scattering can also contribute signi¢cantly to the apparent attenuation (e.g. Richards and Menke, 1983). Scattering tends to transfer the high-frequency energy of the initial body wave into its coda. We can analyze the attenuation in terms of two independent mechanisms. The e¡ective internal friction is the sum of internal friction due to intrinsic loss and loss due to scattering : 31 31 Q31 T ¼ QI þ QS

ð1Þ

where subscripts T, I and S refer to total, intrinsic and scattering, respectively. Many studies of attenuation have been conducted in volcanic areas. Ho-liu et al. (1988) used relative amplitudes of S and P waves from local earthquakes and inferred attenuation anomalies in the Coso geothermal region and in Imperial Valley, Southern California, with QL for shear waves as small as 20 and 30. Sanders (1993) used a similar approach to map a 5^10 km wide anomaly at 7^10 km depth beneath Long Valley caldera with QL 6 13. Commonly, spectral slopes of the initial bodywave phase are used in order to measure t*, a quantity proportional to the time integral of internal friction, Q31 , along the ray path: t
¼

R

ray dt=Q

ð2Þ

Measurements of t* are often done in conjunction with measurements of coda Q, a quantity predicted to combine the e¡ects of intrinsic loss and scattering in a manner di¡erent from Eq. 1 (e.g. Hoshiba, 1991; Wennerberg, 1993). Combining the two types of measurements is therefore a

potential way to separate the two e¡ects. Fehler et al. (1992) found that in the frequency range between 1 and 8 Hz, lithospheric attenuation in the Kanto region of Japan has stronger intrinsic than scattering attenuation. Characteristic combined attenuation over this region, some 200 km across, was estimated as QL = 300. Canas et al. (1998) found a greater dominance of intrinsic attenuation in the Canary Islands over a region of similar size, and a similar average attenuation over the same frequency interval. Their results imply strong frequency dependence. Del Pezzo et al. (1995) worked with local earthquakes recorded at volcano monitoring networks at Etna and Campi Flegrei. Their results are thus relevant to a much smaller and shallower volume in the upper crust. They ¢nd that at low frequencies (1^4 Hz) intrinsic attenuation dominates over scattering attenuation and report a quality factor QL = 20 at 1 Hz at Etna. At frequencies higher than 10 Hz, on the other hand, they ¢nd stronger scattering e¡ects. Other studies of attenuation, often at higher frequencies, have concluded that scattering attenuation is more important than intrinsic attenuation and often ¢nd that P waves are more strongly attenuated than S waves, QK 6 QL (e.g. Del Pezzo et al., 1996; Bianco et al., 1999). The spectral slope method of measuring t* assumes a speci¢c form for the spectrum of the time function of the earthquake source. In order to properly cancel out the shape of the source spectrum, use is made of some spectral ratio (Teng, 1968). This results in the measurement of a di¡erence in t* for two phases (P or S waves for the same event at separate stations, or S and P waves at the same station) or from a mean. This was the method applied by Romero et al. (1997) to microearthquake data in the Geysers geothermal region in California as well as by Evans and Zucca (1988) and Zucca and Evans (1992), who worked at Medicine Lake and Newberry volcanoes in the Cascade range of western USA. In these studies absolute calibration of the quality factor was dif¢cult, but signi¢cant variation of attenuation was imaged and interpreted in terms of a thermal anomaly within the roots of the volcanoes and geothermal £uids in the shallow crust. Wilcock et al. (1995) modeled the spectral content of their

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controlled explosive sources in their work on the East Paci¢c Rise. They modeled a small zone of QP = 20^50 in the vicinity of the magma reservoir beneath this fast-spreading center. For a comprehensive summary of results from attenuation measurements by a range of methods, at varying frequency, on di¡erent scales and in di¡erent geological settings the reader is referred to Sato and Fehler (1998). We present here a study of the seismic attenuation around Rabaul volcano in Papua New Guinea based on spectral-ratio-slope measurements from broadband records collected in the region as part of the Rabaul Earthquake Location and Caldera Structure (RELACS) project.

2. Geological and volcanological setting Rabaul town lies within the nested calderas of a volcanic complex that is named after it. The vol-

79

cano is located at the northeastern corner of the Gazelle Peninsula at the eastern end of New Britain in Papua New Guinea (see Figs. 1 and 2). It is the easternmost volcano of the New Britain trench, where the Solomon plate subducts underneath the Bismarck plate. The Gazelle Peninsula is mostly built up of late Eocene and late Oligocene volcaniclastics with younger limestone sequences outcropping in parts (Lindley, 1988). The region around Rabaul volcano is blanketed with the volcano’s pyroclastic products to a distance of approximately 20 km. Petrologically, Rabaul volcano is a fairly typical island-arc volcano, producing mostly dacitic tephra and pyroclastics in Plinian eruptions and occasionally smaller volumes of basalts and andesites during Strombolian activity (Wood et al., 1995). The eruption history of Rabaul volcano was compiled by Nairn et al. (1995). Three eruption periods are known from the historical record: 1887, 1937^1943 and 1994^present. The

Fig. 1. Position map, showing the location of Rabaul. The islands of New Britain and New Ireland are found within the central square, which is enlarged in the upper right-hand corner. Rabaul is at the northeastern end of New Britain. The approximate locations of the events used are plotted in the enlarged portion of the ¢gure.

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Fig. 2. Map of the Rabaul area. Site names referred to in the text are speci¢ed, including sites of broadband instruments used in this study, which are marked as open triangles. RELACS array land stations (including stations in the permanent network of the Rabaul Volcano Observatory) are plotted as solid circles.

eruption period between 1937 and 1943 is described by Johnson and Threlfall (1985). In these three eruptions Vulcan and Tavurvur, on opposite sides of Blanche Bay, were the active volcanic vents (see Fig. 2). A number of other vents have been active after the collapse of Blanche Bay (1400 BP) (Nairn et al., 1995) and Greene et al. (1986) identify further submarine volcanic cones in Blanche Bay based on detailed depth sounding and single-channel seismics. Together, the recently active volcanic vents form a crudely circular pattern in the center of the Rabaul caldera complex, which may correspond to the Blanche Bay collapse. Detailed seismic monitoring of Rabaul caldera started in 1968 with the establishment of the Rabaul-harbor network operated by the Rabaul Volcano Observatory (RVO). In the early 1970s local earthquake activity increased within the Rabaul

calderas. The activity remained roughly constant until September 1983, when it suddenly increased dramatically and remained high until July 1985. Activity then decreased to a level similar to that before the crisis period, 1983^1985. Signi¢cant uplift associated with this heightened seismic activity was observed centered in Blanche Bay (see Fig. 2). The seismicity delineated an oblate ring centered beneath Blanche Bay and dipping down outward to a depth of approximately 4 km (Mori and McKee, 1987). The uplift was interpreted as due to the injection of melt into two shallow cavities at about 1 km depth within the caldera (Mori et al., 1988). An equally viable explanation of the uplift is a magma injection into the ring fault delineated by the seismicity (Saunders, 2001). After a period of 10 years of roughly constant seismic activity the Rabaul volcano ¢nally erupted in September 1994 (GVN, 1994). Much devasta-

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tion was caused by the eruption in Rabaul town and surrounding areas (Blong and McKee, 1995). Some petrological and geochemical aspects of the eruption are described by Johnson et al. (1995). Roggensack et al. (1996) discuss some aspects of volatiles during the eruption and argue for a crustal magma reservoir at a depth of 2^9 km based on volatiles in melt inclusions. This is consistent with the depth distribution of seismicity at Rabaul if that is interpreted to occur in the roof (Mori and McKee, 1987) and the depth of a signi¢cant low-velocity anomaly found by Finlayson et al. (2003). The most recent of the collapse structures is only 1400 years old and large (Blanche Bay is approximately 100 km2 in area and 500 m deep where it is deepest from rim to bottom). It is unlikely that such a large body, or larger assuming the reservoir is only partially emptied by eruption, at ca. 5 km depth would have cooled signi¢cantly in this short time (Barrie et al., 1999). We therefore expect a substantial volume of partially molten rock beneath Blanche Bay. The seismicity within the caldera may be caused by pressure changes in this reservoir or enhanced geothermal activity due to injections of melt from below starting in the early 1970s and culminating in 1983^ 1985. Rabaul’s tectonic setting results in a wide distribution of regional seismicity. The New Britain trench provides frequent sources from the southeast through south and southwest to west as well as from directly below. The Manus basin spreading center produces earthquakes to the northwest through north. The Weitin transform fault connects the two just to the east of Rabaul and southward to the sharp bend of the New Britain trench.

3. Estimates of the local quality factor from broadband records 3.1. The broadband data from regional earthquakes We seek to quantify the attenuation structure at Rabaul volcano by measuring di¡erential t* from broadband recordings in the region. Two broad-

81

band seismographs were operated in the region in 1997, one within the Rabaul caldera complex, the other some 25 km away from it (see Fig. 2). In the period between August 1997 and January 1998, Geoscience Australia (formerly the Australian Geological Survey Organisation) conducted the RELACS ¢eld project in cooperation with the Research School of Earth Sciences at the Australian National University and Hokkaido University. The overall project objectives are described by Gudmundsson et al. (1999). Project operations are described by Finlayson et al. (2001). Data handling and management is described by Soames et al. (2000). About 80 portable short-period instruments were deployed within 30 km of Rabaul for the duration of the experiment and occupied about 150 recording sites. Two broadband instruments were operated during the period at sites named Vudal and Kaivuna (see Fig. 2). The sites were equipped with Guralp 40T sensors, sensitive to 50 Hz frequency to 30 s period, and RefTek recorders. Both instruments were operated continuously at 50 samples per second. It is the high dynamic range of the recording, 16 bit, which is important rather than the wide bandwidth. The RELACS short-period instrument array recorded at 12-bit dynamic range that is generally insu⁄cient for detailed spectral analysis. Vudal is the site of an agricultural college. It is located at the edge of the alluvial plain around the mouths of the Kerevat and Vudal rivers approximately 25 km SW from Rabaul. The instrument was deployed in a little-used storage building on the outskirts of the college area with a thick concrete £oor sitting on the alluvium. Cultural noise was markedly higher during the day than at night. The instrument was operational for all of the RELACS observational period with the exception of one 3-week gap. Kaivuna is one of the hotels of Rabaul town. It is located on Mango Avenue, in the southern part of old Rabaul town, which was almost completely ruined in the 1994 eruption and its aftermath. This site lies within the Rabaul caldera complex. The instrument was deployed in a small concrete bunker about 100 m behind the hotel and away from Mango Avenue. Conditions at this site were

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not ideal, cultural noise was high during the day in particular. It proved di⁄cult to maintain the instrument because of a di⁄cult power source and after about 1 month of operation the site was £ooded and recording discontinued. Despite recording problems at the Kaivuna site, a number of useful regional earthquakes were recorded simultaneously by both broadband instruments. 3.2. Estimates of N t* We have selected for analysis regional events within a distance of a few hundred km from Rabaul that are well recorded by both broadband stations (see Fig. 1 and Table 1). We include only events with a well-resolved spectrum in the 0.5^5 Hz range. By well-resolved we mean that the signal-to-noise ratio of the spectrum is well above unity over most of the frequency interval. We compute the Fourier spectrum of the cosinetapered time series following the picked arrival time for P and S waves, extending over 256 samples or just over 5 s. The mean is removed from the time series and it is appended with 256 zeros. We also compute the Fourier spectrum of the 256 samples preceding the picks, which we regard as

noise. The spectra are smoothed over a frequency range of 0.2 Hz. The signal-to-noise ratio is estimated as the ratio of the smoothed amplitude spectra derived from those two time windows. We compute the spectra of both P and S waves. The frequency range over which we have a good signal-to-noise ratio for both P and S waves is, in general, 0.5^5 Hz. Fig. 3 summarizes the measurements made for one event. The events we have used are listed in Table 1. Each event is identi¢ed by a number, which is based on its origin time. The ¢rst three digits represent the number of the day of 1997. The following four are the hour and minute of the origin time. The origin times are according to PDE catalogs published by USGS NEIC. For three events not listed in the PDEs we have used the earliest arrival time at the RELACS network. We list in Table 1 the epicentral distance (v), source depth (Z), and magnitude (mb ) according to the published location. Because of potential uncertainty of PDE locations in this region we have estimated the back azimuth (h) and slowness (reciprocal apparent velocity, Va ) of the initial P wave by ¢tting a plane wave to the ¢rst arrivals at short-period seismographs of the RELACS network around the Kaivuna station. A plane wave

Table 1 Events used and measured parameters for events Event

v (‡)

Z (km)

2672023 2681406 2681844 2701959 2711456 2740412 2740757 2741852 2791509 2800738 2801334

1.5

33

0.2 0.9 1.4 0.4

167 100 100 40

1.9

33

1.0 1.4

114 33

h (‡)

Va (km/s)

vtSP (s)

mb

172 85 203 195 105 80 88 110 274 210 167

8.2 8.3 51.0 9.1 12.2 8.1 8.1 8.4 6.3 12.4 7.4

16.6 8.8 17.5 14.6 20.6 8.0 7.9 21.8 12.2 15.2 18.6

4.2 4.6 3.7 4.1 4.7 4.2 4.5 4.0

v, epicentral distance; Z, source depth; mb , magnitude; h, back azimuth; Va , apparent velocity; vtSP , S^P time. Epicentral distance, source depth and magnitude are according to information in Preliminary Determination of Earthquake parameters (PDE), published by the USGS NEIC, if the event is listed there. Back azimuth and apparent velocity are computed by ¢tting a plane wave to ¢rst arrival times at RELACS array. The di¡erential S^P time is the average of measurements at Kaivuna and Vudal stations. Each event is identi¢ed by a number, which is based on its origin time. The ¢rst three digits represent the number of the day of 1997. The following four are the hour and minute of the origin time. The origin times are according to PDE catalogs. For three events not listed in the PDEs we have used the earliest arrival time at the RELACS network.

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83

Fig. 3. An example of seismograms and spectra for magnitude 4.5 event (2800738 at 100 km distance from Rabaul to the SW (210‡ back azimuth). The seismograms in the lower right-hand corner are labeled with Z for the vertical component and T for the transverse component. Noise spectra are shown in a lighter shade of gray.

is not an appropriate description for the phase fronts for such a short propagation distance and the wavefront is commonly complex. There is too much scatter in the arrival times to resolve a consistent curvature. For the listed events the estimated back azimuth is generally consistent with the predicted back azimuth to within 20‡. Finally, we list in Table 1 the average time di¡erence between the arrival of the P wave and the S wave (vtSP ) at both stations in order to get an idea about the distance to the event. We list in Table 2 time factors derived from the S^P times measured at the two stations. We measured the S^P time from both records for each event. This is an imprecise measurement with an

uncertainty of up to 1 s. In order to obtain a consistent estimate of both di¡erential times we predict their di¡erence based on the back azimuth and apparent velocity, assuming that the medium is a Poisson solid (VP /VS = k3). We then use the mean of the two measurements and shift the estimates of the di¡erential S^P times symmetrically from the mean such that the predicted difference is preserved. The numbers, dK and dV , listed in Table 2, are these estimates of di¡erential S^P time divided by (k331). The purpose of this will become apparent in Section 4. We also list the results of our measurements of Nt* in Table 2. We choose to work with spectral ratios because

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Table 2 Measured Nt*s and the time factors dK and dV for each event Event

dK (s)

dV (s)

Nt*P;KV (s)

Nt*S;KV (s)

Nt*SP;K (s)

Nt*SP;V (s)

2672023 2681406 2681844 2701959 2711456 2740412 2740757 2741852 2791509 2800738 2801334

23.6 10.7 23.9 20.8 27.5 10.0 9.5 28.8 17.9 21.8 26.2

21.7 13.3 23.9 19.1 28.8 11.9 12.1 30.8 15.4 19.7 24.6

0.11 V 0.02 0.06 V 0.02 0.07 V 0.03 0.13 V 0.03 0.05 V 0.02 0.10 V 0.03 0.04 V 0.03 0.05 V 0.03 0.05 V 0.03 0.06 V 0.02 0.14 V 0.03

0.33 V 0.02 0.12 V 0.03 0.18 V 0.05 0.33 V 0.04 0.09 V 0.03 0.09 V 0.07 0.10 V 0.03 0.11 V 0.03 0.10 V 0.05 0.17 V 0.04 0.30 V 0.04

0.39 V 0.04 0.18 V 0.04 0.49 V 0.06 0.28 V 0.03 0.24 V 0.02 0.12 V 0.05 0.13 V 0.04 0.18 V 0.04 0.21 V 0.05 0.22 V 0.03 0.17 V 0.04

0.20 V 0.02 0.12 V 0.03 0.39 V 0.04 0.09 V 0.02 0.20 V 0.03 0.11 V 0.04 0.07 V 0.03 0.11 V 0.03 0.16 V 0.04 0.12 V 0.02 0.09 V 0.03

K in the index represents Kaivuna station, V stands for Vudal. Double indices indicate the spectral ratio in each case.

we have a useful signal-to-noise ratio in a frequency range comparable to predicted corner frequencies for the magnitudes of the available events. We therefore need a measurement procedure which cancels the e¡ects of the source. For each event we have two wave types, P and S, at two stations, K and V. Here we use K to represent the Kaivuna station, inside the caldera, and V to represent the Vudal station, outside the caldera. We can estimate four di¡erent spectra and from those four sensible spectral ratios: the spectral ratio of the two P waves at the two stations; the spectral ratio of the two S waves at the two stations; the spectral ratio of S and P waves at the Kaivuna station; the spectral ratio of S and P waves at the Vudal station. The remaining two possibilities are not sensible and in fact the four above only yield three independent pieces of information. We nevertheless measure the above four and index the resulting di¡erential t* estimates from the spectral ratio slopes in the following manner : The subscript contains the wave type ¢rst, the station second, the two separated by a comma. If the wave type is repeated, the estimate is from a spectral ratio of the two wave types at a ¢xed station. If the station index is repeated the estimate is from a spectral ratio of the same wave type at the two stations. We normalize spectra from Kaivuna by spectra from Vudal and Swave spectra by P-wave spectra. Thus, Nt*P;KV is derived from the slope of the spectral ratio of Kaivuna P waves and Vudal P waves. Nt*SP;K is

derived from the slope of the spectral ratio of S waves and P waves at Kaivuna. The Nt* is, in each case, estimated over the frequency range from 0.5 to 5 Hz by a weighted least-squares ¢t to a line with each ratio weighted by a function which depends on the signal-to-noise ratios. This weighting function can be represented by: w ¼ clipð0; 1; minflog10 ðas =an Þ; log10 ðbs =bn ÞgÞ

ð3Þ

Here clip(0,1; x) is the value of x, unless x exceeds 1 or is less than 0. In those cases it is clipped, i.e. ¢xed at 1 or 0, respectively. The function min{x,y} is the smaller of x and y. The spectral ratio is a/b. The subscripts s and n refer to signal and noise, respectively. Thus, if the signal-to-noise ratio exceeds 10 for both amplitude spectra the weight is 1. If the signal-to-noise ratio is less than 1 for either spectrum the weight is 0. The uncertainty estimate of each t* estimate is the formal regression error from the linear ¢t to each spectral ratio. It is stated in Table 2 as one standard deviation. A spectral ratio of an S wave normalized by a P wave only makes sense if we can assume that the source spectrum is the same for the two wave types. This is perhaps not strictly true. See e.g. Gudmundsson et al. (1994) for a discussion of potential e¡ects of unequal source spectra for two wave types. Neither site, Vudal or Kaivuna, was deployed on bedrock. Both sites could therefore be signi¢cantly a¡ected by localized site ampli¢cation, which may be frequency-dependent and thus af-

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fect our measurements of Nt*. In order to evaluate this e¡ect we have utilized the receiver-functiontype method of Nakamura (1989), which approximates the site response to S waves with the spectral ratio of the horizontal and the vertical components of motion. The advantage of this method is that it does not require a station deployed on bedrock for reference and is not dependent on assumptions about the relative attenuation structure beneath the site to be evaluated and the reference site. The disadvantage is that a reliable estimate requires averaging over a number of events (Huang and Teng, 1999). Furthermore, the method has not been developed for P waves. We compute the average spectral ratio of the transverse and vertical components of motion for the S waves for all the 11 events used in this study and at both sites. The resulting estimates of site response are weakly frequency-dependent. At Vudal the ampli¢cation decays smoothly with frequency with an average slope equivalent to t* = 0.10 s. Kaivuna possesses a resonance peak at about 2 Hz with an ampli¢cation by a factor of 2 over adjacent frequencies but otherwise the site response decays with increasing frequency. The systematic e¡ect on estimates of t* is 0.07 s. The combined e¡ect on our measurements of Nt*S;KV is the di¡erence between the overall spectral slopes of the site response estimates, or 0.03 s. We can only assume that the e¡ect is the same for P waves. Because the e¡ect is marginal, i.e. comparable to error estimates in Table 2, we choose not to correct for it, but keep in mind that the estimates of Nt*X ;KV (X = P or S) in Table 2 may be slightly biased to the low side (by 0.03 s). The results in Table 2 have some striking features worth highlighting. All the estimates of differential t* are positive. In other words, the attenuation at Kaivuna station is always greater than at Vudal station and the attenuation of S waves is always greater than the attenuation of P waves. Furthermore, the S/P spectral-ratio slope at Kaivuna is always larger than the corresponding slope at Vudal. These generalizations are in some cases marginal, i.e. measurements or di¡erences between them are comparable to uncertainty estimates. Taken as a whole the data do, however, clearly indicate that attenuation is stron-

85

ger around and beneath Kaivuna, or the caldera region, than at Vudal. 3.3. The frequency dependence of attenuation The above analysis assumes that Q is not frequency-dependent. However, both laboratory experiments and many ¢eld observations suggest that Q is frequency-dependent. Q generally increases with increasing frequency. Intrinsic attenuation Q increases as a power law with frequency in many laboratory measurements (Kampfmann and Berckhemer, 1985; Jackson et al., 1992): Q ¼ Q 0 ð X =X 0 Þ n

ð4Þ

(X is frequency, X0 is reference frequency and Q0 the quality factor at that reference frequency) with a power-law exponent of n = 0.1^0.3. Apparent Q increases even faster with frequency according to a number of ¢eld experiments. We note that if the power-law exponent of the frequency dependence were n = 0.2, then Q would vary by about 40% over the frequency band from 0.5 to 5 Hz. It is evident from Fig. 3 that we cannot resolve a curvature on the spectral ratios from which we measure the spectral slope. In order to examine the potential frequency dependence further we are forced to stack the data to reduce scatter in the spectral ratios. To compute a meaningful stack we assume that the frequency dependence is the same in all cases. Each log-spectral ratio is assumed to have the form : S j ð i Þ ¼ aj 3ZX

R

dt=Qj

ð5Þ

where Qj = Q0;j (X/X0 )n (aj stands for the absolute amplitude and j indexes the spectral ratio) and it is ¢t with a Q that is not frequency-dependent. The residual log-spectral ratio is then approximately : R vSi ðX Þ ¼ ZX ð13ðX 0 =X Þn Þ dt=Q0;i ¼ ZX ð13ðX 0 =X Þn Þt
i

ð6Þ

If we stack the residual log-spectral ratios, normalized by ZXti *, we recover the frequency dependence:

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GvS i ðX Þ=ZX t
i f ¼ 13ðX 0 =X Þn

ð7Þ

Here Gf represents an ensemble average. This normalization by the slope should cancel out di¡erential e¡ects of varying path length or variable attenuation. The left-hand side of Eq. 7 is plotted in Fig. 4 together with a one standard deviation error bound and the right-hand side predicted for n = 30.1, 0.1, and 0.3. The reference frequency is taken to be the center of the frequency band that we use, X0 = 2.75 Hz. The indication from Fig. 4 is that we cannot resolve a frequency dependence and that if the frequency dependence is the same power law for all spectral ratios then its exponent is less than n = 0.1. If anything, the indication is that the exponent is slightly negative, i.e. that attenuation increases with frequency (Q reduces), which may indicate that the contribution of scattering to the apparent attenuation increases with frequency. 3.4. The ratio Nt*S /Nt*P and intrinsic versus scattering attenuation The estimates of di¡erential t* in Table 2, Nt*S;KV and Nt*P;KV , are plotted against each other in Fig. 5. In light of the results in Section 3.2 we interpret these ratios as manifestations of a local attenuation anomaly in the vicinity of the Kaivuna station, inside the Rabaul caldera complex.

Fig. 4. The apparent frequency dependence of attenuation in the used frequency range from 0.5 to 5.0 Hz. The thick solid line represents the estimate from the data with one standard deviation error bounds (thin solid lines). The dashed lines show predictions of power-law behavior of attenuation with indicated power-law exponents.

Fig. 5. The di¡erential t* for S waves versus P waves. Both are derived from the spectral measurements at Kaivuna normalized by spectra from Vudal.

The best line (data weighted by errors) through the origin and through the points in Fig. 5 has a slope of 2.6. This implies, if the medium has a velocity ratio VP /VS = 1.8^2.0, that the attenuation ratio is QK /QL = 1.3^1.5 for P waves versus S waves. This is a fairly high ratio, although not as high as one would expect for purely intrinsic attenuation due to creep mechanisms (Anderson, 1989). If attenuation is entirely due to a viscous component of the shear modulus we would expect QK /QL = 2.2^2.6, depending on the Poisson ratio. In contrast, many ¢eld studies, which conclude that scattering attenuation is the dominant mechanism of attenuation, ¢nd QK /QL 6 1. This suggests that both scattering and internal friction contribute to the attenuation. However, scattering attenuation is generally strongly frequency-dependent. The fact that we see little evidence for frequency dependence may suggest that the local attenuation anomaly beneath the Rabaul caldera complex is more strongly a¡ected by intrinsic attenuation.

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87

4. Analysis of Nt*s in terms of local and regional attenuation The Nt* derived from the spectral ratios of either wave type at the two stations (Nt*S;KV and Nt*P;KV ) are in all cases positive. This indicates that attenuation is stronger in the vicinity of the Kaivuna site inside the Rabaul caldera complex. Furthermore, the Nt* derived from the spectral ratios of the two wave types at the Kaivuna station (Nt*SP;K ) is always bigger than at the Vudal station (Nt*P;KV ), again indicating stronger attenuation inside the caldera than outside. We cannot de¢nitively resolve the anomaly to be con¢ned to Rabaul caldera or the immediate vicinity of the volcano. However, the simplest possible interpretation is in terms of a single anomaly near Kaivuna. Moreover, it makes sense geologically to associate the anomaly with the volcano, in particular the caldera (fractured rock, geothermal £uids, magma reservoir). Thus we formulate an interpretation of the observations in terms of a local e¡ect in the vicinity of Kaivuna superimposed on a regional background. We argue that the attenuation structure of the region is characterized by a regional attenuation structure, which may depend on the direction from Rabaul volcano, but is dominated by a local anomaly of high attenuation at Rabaul volcano and at shallow depths. We generalize this hypothesis in Fig. 6. A regional earthquake occurs at point P. It is observed at points K and V, adjacent to the localized anomaly and distal to it, respectively. We can then characterize the slope of the spectra of P and S waves to the two recording sites by: t
P;V ¼ tP;V =QK

the shaded area of Fig. 6), tP;V , tS;V , tP;K , tS;K are the travel times from the source to stations V and K for P and S waves, and dP and dS are the travel times spent by P and S waves, respectively, within the localized attenuation anomaly. These expressions ignore the source spectrum, instrument response, geometrical spreading e¡ects and other factors that do not a¡ect the slope of the spectral ratio if the source is the same for P and S waves. Neither do they include frequencydependent propagation e¡ects, such as local resonances beneath each station; however, those are small (see Section 3.2). The spectral ratios then become : Nt
P;KV ¼ ðtP;K 3tP;V 3d P Þ=QK þ d P =qK

ð9aÞ

Nt
S;KV ¼ ðtS;K 3tS;V 3d S Þ=QL þ d S =qL

ð9bÞ

Nt
SP;V ¼ tS;V =QL 3tP;V =QK

ð9cÞ

Nt
SP;K ¼ ðtS;K 3d S Þ=QL þ d S =qL 3ðtP;K 3d P Þ=QK 3d P =qK ð9dÞ

t
S;V ¼ tS;V =QL t
P;K ¼ ðtP;K 3d P Þ=QK þ d P =qK t
S;K ¼ ðtS;K 3d S Þ=QL þ d S =qL

Fig. 6. A schematic ray diagram distinguishing between separation of regional attenuation structure and a local anomaly near Kaivuna station.

ð8Þ

where QK and QL are the path-averaged regional quality factors (along paths outside the shaded area in Fig. 6), qK and qL are the path-averaged quality factors within the local anomaly (inside

We assume that the S-wave paths are equivalent to the P-wave paths and that for the purpose of predicting the regional time terms the medium is a Poisson solid, i.e. VP /VS = k3, and de¢ne: d K ¼ ðtS;K 3tP;K Þ=ðk331Þ d V ¼ ðtS;V 3tP;V Þ=ðk331Þ

ð10Þ

We then combine the quality factors into ratios:

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Q ¼ QK =QL ; R ¼ V P =V S WqK =qL

ð11Þ

The above expressions for the spectral ratios (Eq. 9) simplify to: Nt
P;KV ¼ ðd K 3d V 3d P Þ=QK þ d
P

ð12aÞ

Nt
S;KV ¼ k3 Q ðd K 3d V 3d P Þ=QK þ R d
P

ð12bÞ

Nt
SP;V ¼ ðk3 Q 31Þd V =QK

ð12cÞ

Nt
SP;K ¼ ðk3 Q 31Þðd K 3d P Þ=QK þ ð R 31Þd
P

ð12dÞ

where d*P = dP /qK . Only three of these equations are independent. The redundancy of Eq. 12 can be expressed as: Nt
S;KV 3Nt
P;KV ¼ Nt
SP;K 3Nt
SP;V

ð13Þ

The estimates in Table 2 are consistent with Eq. 13 to well within uncertainty for all 11 events. We can obviously not resolve the ¢ve unknowns, dP , QK , Q, d*P , and R, from the three independent equations for each event. We argue that the parameter R should be approximately the same for all events, because the rays into Kaivuna are all sampling the same local attenuation anomaly, albeit with varying geometry. We therefore use the estimate in Section 3.4 derived from Fig. 5, R = 2.7. Note, however, that Eq. 12 demonstrates how the apparent constant of proportionality between Nt*S;KV and Nt*P;KV is not a pure measure of R. Similarly, we assume that Q is the same for all regional paths. The remaining three parameters are QK , d*P , and tP . QK and d*P relate directly to attenuation, regional and local, respectively. dP , on the other hand, is the time spent by each ray into Kaivuna station in the local anomaly. This is poorly constrained, but if the anomaly is con¢ned to the caldera complex we would expect dP 6 3 s. The regional attenuation can be estimated from Eq. 12c and information in Table 2. The time constant, dV , is in the range 10^30 s, while Nt*SP;V is 0.1^0.2 s, except for event 2681844. Thus, if we assume Q is of order 2 we estimate QK V350. We can then estimate the regional term in Eq. 12a,b.

While Nt*P;KV lies in the range 0.05^0.15 s, the regional term in Eq. 12a is (dK 3dV 3dP )/QK 6 0.015. While Nt*S;KV lies in the range 0.10^0.35 s, the regional term in Eq. 12b is k3Q (dK 3dV 3dP )/ QK 6 0.05. In other words, the right-hand sides of Eq. 12a,b are dominated by the local term, at least for those events where Nt*P;KV and Nt*S;KV are measured relatively large. Since the local term dominates Eq. 12a,b we can combine the two to establish a reasonable estimate of R, as was done in Section 3.4 and Fig. 5. Since the regional term is smaller than the local term in all cases only Eq. 12c constrains Q. However, Q occurs in Eq. 12c in the ratio Q /QK . Therefore, we have no information to separate Q and QK , we can only constrain their ratio. dP occurs as an uncertain quantity in minor terms. We cannot constrain dP and are forced to simply guess its value. We proceed to estimate QK for individual events using Eq. 12c and assuming Q = 2. The estimate of QK will scale directly with the choice of Q. We then proceed to estimate d*P from Eq. 12a,b by ¢rst correcting them by an estimate of the regional term based on the individual QK estimate and a guess of dP . Since Kaivuna station is in the northern part of the caldera we assign a value of dP = 1.5 s for events from the east, north and west and dP = 3 s for events from the south: d
P ¼ Nt
P;KV 3ðd K 3d V 3d P Þ=QK and

R d
P ¼ Nt
S;KV 3k3 Q ðd K 3d V 3d P Þ=QK

ð14Þ

Thus, we re-evaluate R by the same regression analysis as was applied in Section 3.4, use the best estimate of R to normalize the second estimate of d*P in Eq. 14, and ¢nally average the two estimates from Eq. 14. The best estimate of R is: R = 2.7 V 0.5. The results for QK and d*P are presented in Table 3. The estimates of R and d*P depend weakly on the guessed value of Q. Changing Q to 1.5 reduces the estimates of QP by 35%, does not a¡ect the estimate for R and a¡ects estimates of d*P by less than 10%, i.e. insigni¢cantly. Error estimates include all sources of error other than in the parameters Q and R since those

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O. Gudmundsson et al. / Journal of Volcanology and Geothermal Research 130 (2004) 77^92 Table 3 Estimated regional quality factor for QP , and local t* for compressional waves, d*P Event

QP

dP * (s)

2672023 2681406 2681844 2701959 2711456 2740412 2740757 2741852 2791509 2800738 2801334

267 V 39 273 V 88 151 V 22 523 V 144 354 V 65 266 V 119 426 V 218 690 V 211 237 V 75 404 V 88 673 V 252

0.13 V 0.01 0.07 V 0.02 0.09 V 0.02 0.13 V 0.01 0.06 V 0.01 0.08 V 0.03 0.05 V 0.02 0.05 V 0.02 0.05 V 0.02 0.06 V 0.01 0.13 V 0.02

are systematic. Uncertainties of the quantities in Table 3 range from 10% to 50%. The estimates in Table 3 may be translated to parameters for S waves by: d*S = Rd*P , and QS = QP / Q .

5. Discussion The estimates in Table 3, translated to S-wave parameters, have been plotted against back azimuth in Fig. 7. There is no suggestion of a correlation between either parameter QL or d*S and epicentral distance.

Fig. 7. Regional internal friction (top) and the local d*S within Rabaul caldera (bottom) as functions of back azimuth from Kaivuna station.

89

The estimates of regional attenuation are scattered around 1/QL V0.050 with a variation of about 0.015. It is of no surprise that this parameter varies, because the paths are in fact very different. The order of magnitude is, however, consistent. One point is distinct from the rest, i.e. at approximately 205‡ back azimuth. This point is for event number 2681844, which focus was nearly directly beneath Rabaul at a depth of 167 km. Note also the very high apparent velocity for this event in Table 1, i.e. apparent steep angle of incidence. The steep path for this event is quite di¡erent from those for all the other events, which refract in the upper lithosphere and crust. The average attenuation along the regional paths is about QL = 180. The path-averaged attenuation along the vertical path for event 2681844 is QL = 75. This suggests that attenuation may be greater in the lithosphere around the volcano. Again, these values scale with an arbitrary choice of Q and are estimated assuming Q = 2. The local attenuation depends on back azimuth from the Kaivuna station. It is scattered between d*S = 0.1 and 0.2 s (accounting for the e¡ect of site ampli¢cation would increase this value insigni¢cantly) except for paths from the south where it is twice as high. This lends further support to association of the local attenuation anomaly with the Rabaul caldera complex. Kaivuna is in the

Fig. 8. Local d*S variation with orientation of rays into Kaivuna. Estimates have been averaged within clusters. Outline of Rabaul caldera complex is drawn with thick line for reference as are recently active eruptive vents, here marked as solid triangles.

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northern part of the caldera complex and north of the most recent collapse structure, the region of concentrated seismicity and region of recently active volcanic vents. The azimuthal variation of d*S is shown in map view in Fig. 8. There it becomes clear that those paths, which pass beneath the region bounded by recently active volcanic vents, have higher attenuation than those that do not. We cannot con¢ne this anomaly, but if it is con¢ned to the caldera, such that path lengths through it into Kaivuna are on the order of 10^ 15 km, then we would estimate an overall quality factor for the anomaly of qL = 15^20. This indicates a high level of attenuation, similar to estimates in a similar frequency range at other volcanos (e.g. Ho-liu et al., 1988; Sanders, 1993; Del Pezzo et al., 1995; Wilcock et al., 1995) and marginally higher than measured in the laboratory for volcanic rocks near their solidus (Kampfmann and Berckhemer, 1985). It is possible that melt is present along a portion of some of the paths measured here, but only in small volume fractions. A more likely, general explanation of high attenuation in the Rabaul caldera region is fracturing of the caldera rocks and presence of geothermal £uids there as well as scattering effects. It is interesting that two of the observed S waves at Kaivuna station arrive at a back azimuth pointing under the caldera and at an apparent velocity corresponding to crustal or uppermost mantle velocity. Using the velocity model of Finlayson et al. (2003) for average caldera structure the ray paths for these events are expected to pass under the central south caldera at depths around 5 km or greater. This, at a glance, indicates that there is no shear-wave shadow corresponding to the postulated magma reservoir beneath Blanche Bay (Mori and McKee, 1987; Roggensack et al., 1996; Finlayson et al., 2003). Note also that the velocities imaged by Finlayson et al. (2003) ( s 4 km/s) at 3^5 km depth beneath the central caldera are not so low as to require a molten body there. However, this may be a smoothing e¡ect due to lack of resolution. Furthermore, their low-velocity anomaly terminates at a depth of about 6 km and it is possible that the shear waves observed here from regional events have passed underneath it.

6. Conclusions Measurements of di¡erential t* between S and P waves and at two separate recording sites in the Rabaul region, inside and distal to Rabaul caldera, have led to estimates of t* due to a localized attenuation anomaly associated with Rabaul caldera. This indicator of attenuation depends on back azimuth from Kaivuna station in the northern part of the caldera. Attenuation appears to be strongest to the south of Kaivuna. The apparent ratio of S- and P-wave attenuation is qK /qL W1.3^1.5. No frequency dependence is resolved in the band between 0.5 and 5 Hz. We argue that the measured attenuation is strongly a¡ected by intrinsic attenuation, and that the low Q of the caldera region is caused by fracturing of the caldera rocks, presence of geothermal £uids there, as well as scattering e¡ects.

Acknowledgements We thank the Australian Agency for International Development (AusAID) for their funding of the RELACS program under which the broadband recordings were undertaken. The Australian National University supplied the instruments used at the two broadband sites. Sta¡ at the Rabaul Volcano Observatory, Geoscience Australia (formerly Australian Geological Survey Organisation), Australian National University and University of Wisconsin as well as RELACS program sta¡ provided valuable help in the ¢eld. In particular, we would like to thank Jonathon Kudua, Lee Powell, Geo¡ Clitheroe, Heather Miller, Lucinda Sawyer, and Armando Arcidiago. The manuscript has bene¢ted from the constructive reviews of two anonymous reviewers. We acknowledge the Danish Research Foundation (Grundforskningsfonden) for supporting O.G. in the analysis of the work presented here.

References Anderson, D.L., 1989. Theory of the Earth. Blackwell Sci. Publ., Boston, MA, 366 pp.

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O. Gudmundsson et al. / Journal of Volcanology and Geothermal Research 130 (2004) 77^92 Barrie, C.T., Cathles, L.M., Erendi, A., Schweiger, H., Murrey, C., 1999. Heat and £uid £ow in volcanic-associated, massive, sul¢de-forming hydrothermal systems. Rev. Econ. Geol. 8, 201^219. Bianco, F., Castellano, M., Del Pezzo, E., Ibanez, J.M., 1999. Attenuation of short-period seismic waves at Mt Vesuvius, Italy. Geophys. J. Int. 138, 67^76. Blong, R., McKee, C., 1995. The Rabaul Eruption 1994: Destruction of a Town. Natural Hazard Research Centre, Macquarie University, Sydney, 52 pp. Canas, J.A., Ugalde, A., Pujades, L.G., Carracedo, J.C., Soler, V., Blanco, M.J., 1998. Intrinsic and scattering seismic attenuation in the Canary Islands. J. Geophys. Res. 103, 15037^15050. Del Pezzo, E., Ibanez, J., Morales, J., Akinci, A., Maresca, R., 1995. Measurements of intrinsic and scattering seismic attenuation in the crust. Bull. Seism. Soc. Am. 85, 1373^1380. Del Pezzo, E., Simini, M., Ibanez, J.M., 1996. Separation of intrinsic and scattering Q for volcanic areas: a comparison between Etna and Campi Flegrei. J. Volcanol. Geotherm. Res. 70, 213^219. Evans, J.R., Zucca, J.J., 1988. Active, high-resolution seismic tomography of compressional-wave velocity and attenuation structure at Medicine Lake volcano, northern California Cascade Range. J. Geophys. Res. 93, 15016^15036. Fehler, M., Hoshiba, M., Sato, H., Obara, K., 1992. Separation of scattering and intrinsic attenuation for the KantoTokai region, Japan, using measurements of S-wave energy versus hypocentral distance. Geophys. J. Int. 108, 787^ 800. Finlayson, D.M., Gudmundsson, O., Itikarai, I., Saunders, S., Powell, L., Thurber, C.H., Shimamura, H., Nishimura, Y., 2001. The Rabaul Earthquake Location and Caldera structure (RELACS) Program: Operations report. Australian Geological Survey Organisation Record 2001/01, Canberra, 67 pp. Finlayson, D.M., Gudmundsson, O., Itikarai, I., Nishimura, Y., Shimamura, H., Johnson, W.R., 2003. Rabaul Volcano, Papua New Guinea: island arc crustal architecture and seismic tomographic imaging of an active volcanic caldera. J. Volcanol. Geotherm. Res. 124, 153^171. Greene, H.G., Ti⁄n, D.L., McKee, C.O., 1986. Structural deformation and sedimentation in an active caldera, Rabaul, Papua New Guinea. J. Volcanol. Geotherm. Res. 30, 327^ 356. Gudmundsson, O., Kennett, B.L.N., Goody, A., 1994. Broadband observations of upper-mantle seismic phases in northern Australia and the attenuation structure in the upper mantle. Phys. Earth Planet. Inter. 84, 207^226. Gudmundsson, O., Johnson, W.R., Finlayson, D.M., Nishimura, Y., Shimamura, H., Terashima, A., Itikarai, I., Thurber, C., 1999. Multinational seismic investigation focuses on Rabaul Volcano. EOS Trans. AGU 80, 269^273. GVN, 1994. Rabaul. Bull. Global Volcanism Network 19(8), 2^6, and 19(9), 4^7. Ho-liu, P., Kanamori, H., Clayton, R.W., 1988. Applications of attenuation tomography to Imperial Valley and Coso-In-

91

dian Wells region, Southern California. J. Geophys. Res. 93, 10501^10520. Hoshiba, M., 1991. Simulation of multiple scattered codawave excitation based on the energy conservation law. Phys. Earth Planet. Inter. 67, 123^136. Huang, H.-C., Teng, T.-L., 1999. An evaluation on H/V ratio vs. spectral ratio for site-response estimation using the 1994 Northridge earthquake sequance. Pure Appl. Geophys. 156, 631^649. Jackson, I., Paterson, M.S., FitzGerald, J.D., 1992. Seismic Y heim dunite: an experwave dispersion and attenuation in A imental study. Geophys. J. Int. 108, 517^534. Johnson, W.R., Threlfall, N.A., 1985. Vocano Town - The 1939^43 Eruptions at Rabaul. Robert Brown, Bathurst, 151 pp. Johnson, W.R., McKee, C.O., Eggins, S., Woodhead, J., Arculus, R.J., Chappel, B.W., 1995. Taking petrologic pathways toward understanding Rabaul’s restless caldera. EOS Trans. AGU 76, 171. Kampfmann, W., Berckhemer, H., 1985. High-temperature experiments on the elastic and anelastic behaviour of magmatic rocks. Phys. Earth Planet. Inter. 40, 223^247. Lindley, D., 1988. Early Cainozoic stratigraphy and structure of the Gazelle Peninsula, east New Britain: An example of extensional tectonics in the New Britain arc-trench complex. Aust. J. Earth Sci. 35, 231^244. Mori, J., McKee, C., 1987. Outward-dipping ring-fault structure at Rabaul caldera as shown by earthquake locations. Science 235, 193^195. Mori, J., McKee, C., Itikarai, I., Lowenstein, P., de Saint Ours, P., Talai, B., 1988. Earthquakes of the Rabaul seismo-deformation crisis, September 1983-July 1985: seismicity on a caldera ring fault. In: Latter, J.H. (Ed.), Volcanic Hazard Assessment and Monitoring. Springer, New York, pp. 429^462. Nairn, I.A., McKee, C.O., Talai, B., Wood, C.P., 1995. Geology and eruptive history of the Rabaul caldera area, Papua New Guinea. J. Volcanol. Geotherm. Res. 69, 255^284. Nakamura, Y., 1989. A method for dynamic characteristics estimation of subsurface using microtremor on the ground surface. Q. Rept. Railway Techn. Res. Inst. 30, 25^33. Richards, P.G., Menke, W., 1983. The apparent attenuation of a scattering medium. Bull. Seism. Soc. Am. 73, 1005^1021. Roggensack, K., Williams, S.N., Schaefer, S.J., Parnell, R.A., 1996. Volatiles from the 1994 eruptions of Rabaul: Understanding large caldera systems. Science 273, 490^493. Romero, A.E., McEvilly, T.V., Majer, E.L., 1997. 3D microearthquake attenuation tomography at the Northwest Geysers geothermal region, California. Geophysics 62, 149^167. Sanders, C.O., 1993. Reanalysis of S-to-P amplitude ratios for gross attenuation structure, Long Valley Caldera, California. J. Geophys. Res. 98, 22069^22079. Sato, H., Fehler, M.C., 1998. Seismic wave propagation and scattering in the heterogeneous Earth. AIP Series in Modern Acoustics and Signal Processing. Springer, New York, 308 pp. Saunders, S.J., 2001. The shallow plumbing system of Rabaul

VOLGEO 2705 29-12-03

92

O. Gudmundsson et al. / Journal of Volcanology and Geothermal Research 130 (2004) 77^92

caldera: a partially intruded ring fault? Bull. Volcanol. 63, 406^420. Soames, C.L., Gudmundsson, O., Finlayson, D.M., 2000. The Rabaul Earthquake Location and Caldera structure (RELACS) Program: post survey data processing. Australian Geological Survey Organisation Record 2000/40, Canberra, 19 pp. Teng, T.L., 1968. Attenuation of body waves and the Q structure of the mantle. J. Geophys. Res. 73, 2195^2208. Wennerberg, L., 1993. Multiple-scattering interpretations of coda-Q measurements. Bull. Seism. Soc. Am. 83, 279^ 290.

Wilcock, W.S.D., Solomon, S.C., Purdy, G.M., Toomey, D.R., 1995. Seismic attenuation structure of the East Paci¢c Rise near 9‡30PN. J. Geophys. Res. 100, 24147^24165. Winkler, K.W., Nur, A., 1982. Seismic attenuation e¡ects of pore £uids and frictional sliding. Geophysics 47, 1^15. Wood, C.P., Nairn, I.A., McKee, C.O., Talai, B., 1995. Petrology of the Rabaul caldera area, Papua New Guinea. J. Volcanol. Geotherm. Res. 69, 285^302. Zucca, J.J., Evans, J.R., 1992. Active, high-resolution compressional wave attenuation tomography at Newberry volcano, Central Cascade Range. J. Geophys. Res. 97, 11047^ 11055.

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