Bulletin of the Seismological Society of America, Vol. 98, No. 5, pp. 2428–2448, October 2008, doi: 10.1785/0120070213
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Three-Dimensional P-Wave Velocity Structure and Precise Earthquake Relocation at Great Sitkin Volcano, Alaska by Jeremy D. Pesicek, Clifford H. Thurber, Heather R. DeShon,* Stephanie G. Prejean, and Haijiang Zhang†
Abstract
Waveform cross-correlation with bispectrum verification is combined with double-difference tomography to increase the precision of earthquake locations and constrain regional 3D P-wave velocity heterogeneity at Great Sitkin volcano, Alaska. From 1999 through 2005, the Alaska Volcano Observatory (AVO) recorded ∼1700 earthquakes in the vicinity of Great Sitkin, including two ML 4.3 earthquakes that are among the largest events in the AVO catalog. The majority of earthquakes occurred during 2002 and formed two temporally and spatially separate event sequences. The first sequence began on 17 March 2002 and was centered ∼20 km west of the volcano. The second sequence occurred on the southeast flank of Great Sitkin and began 28 May 2002. It was preceded by two episodes of volcanic tremor. Earthquake relocations of this activity on the southeast flank define a vertical planar feature oriented radially from the summit and in the direction of the assumed regional maximum compressive stress due to convergence along the Alaska subduction zone. This swarm may have been caused or accompanied by the emplacement of a dike. Relocations of the mainshock–aftershock sequence occurring west of Great Sitkin are consistent with rupture on a strike-slip fault. Tomographic images support the presence of a vertically dipping fault striking parallel to the direction of convergence in this region. The remaining catalog hypocenters relocate along discrete features beneath the volcano summit; here, low P-wave velocities possibly indicate the presence of magma beneath the volcano. Online Material: Checkerboard resolution tests and mapview model results.
Introduction Great Sitkin Island is part of the Andreanof Island chain in the western Aleutian arc and is located approximately 35 km northeast of the village and former naval station of Adak (Fig. 1). Great Sitkin volcano, a stratovolcano composed of alternating layers of andesite and basalt, forms the bulk of the island. Great Sitkin has erupted at least eight times since 1792. The most recent eruption in 1974 generated minor ash falls and emplaced a lava dome in the summit crater (Miller et al., 1998; Waythomas et al., 2003; Power et al., 2004). Great Sitkin is seismically active, with active fumaroles and degassing occurring at the summit and on the south flank (Waythomas et al., 2003). In 1999, the Alaska Volcano Observatory (AVO) deployed a telemetered * Present address: Center for Earthquake Research and Information, University of Memphis, 3890 Central Avenue, Memphis, Tennessee 38152. † Present address: Department of Earth, Atmospheric and Planetary Sciences, Massachusetts Institute of Technology, 77 Massachusetts Avenue, Cambridge, Massachusetts 02139.
six station short-period seismic network around the volcano (Fig. 1). This network and neighboring Andreanof Island stations continuously monitor seismicity at Great Sitkin and nearby volcanoes to accurately assess changes in volcanic hazard. Most events at Great Sitkin are identified as volcano-tectonic (VT) earthquakes, which are indicative of brittle failure (McNutt, 2005). A limited number of longperiod (LP) and tremor events, which are associated with fluid processes, have also been recorded (Moran et al., 2002; Power et al., 2004). In 2002, the seismicity rate at Great Sitkin temporally increased when two spatially and temporally separate, off-summit earthquake sequences occurred (Fig. 2; Moran et al., 2002). The AVO catalog locations of events at Great Sitkin are of limited value for detailed studies of volcanic processes, however, due to the large errors and poorly constrained uncertainties. Event location accuracy depends on several factors, each of which presents its own inherent difficulties. In volcanic regions, noisy signals and sparse networks are com-
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Figure 1. Regional map of Great Sitkin volcano and vicinity showing local and regional seismic stations (triangles), volcanic summit (diamond), and 1 arc-second Shuttle Radar Topography Mission (SRTM) topography. Stations ADK and ADAG are located on Adak Island. Also shown are the present volcanic crater rim (gray line around the diamond) and active fumarole (gray star) from Waythomas et al. (2003). Inset map shows the position of Great Sitkin along Alaska’s Aleutian arc. monplace, and catalog locations are frequently calculated using 1D velocity models. The small network sizes and unusual geometries imposed by deployment on steep and poorly consolidated slopes limit ray path coverage. Unmodeled velocity structure affects travel-time calculations. Inaccuracies in arrival time picks for primary seismic phases also map into location precision and accuracy. When solving for event location and velocity structure jointly, these errors propagate through to the velocity model, blurring the resulting velocity image. In this study, we attempt to decrease errors due to inaccuracies in arrival time picks and unmodeled velocity structure from the Great Sitkin seismic data. First, we employ the waveform-alignment technique of Du, Thurber, and Eberhart-Phillips (2004), which uses the bispectrum (BS) method to verify cross-correlation-derived differential times for groups of similar earthquakes at a common station. The BS method correlates waveforms in the third-order spectral domain, significantly decreasing errors due to correlated noise inherent in many signals and providing a statistically independent estimate for waveform alignment and, therefore, increasing the reliability of differential times obtained by
cross correlation (CC). Second, we use the double-difference (DD) tomography technique of Zhang and Thurber (2003) to jointly relocate hypocenters and determine the 3D velocity structure in the region using absolute and more accurate differential times. By using BS-verified CC data to solve for the DD relocations and velocity model, we substantially increase the precision of each (e.g., Du, Thurber, Reyners, et al., 2004; Zhang and Thurber, 2006). These results allow the fine-scale structure of seismicity to be determined and allow us to gain greater insight into the tectonic and magmatic processes at work beneath Great Sitkin and the surrounding region. Hypocenter relocations reveal a hydrothermally active summit conduit and two active faults, one of which may be a developing regional crustal fault while the other may have been caused or accompanied by magma injection. The velocity model indicates low-velocity regions beneath the summit and anomalous velocities associated with the areas of offsummit activity that occurred in 2002.
Data The data for this study are composed of waveform time series, arrival times, and catalog locations for more
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Figure 2.
AVO catalog seismicity for Great Sitkin volcano from September 1999 through December 2005. Hypocenters are located using the HYPOELLIPSE location algorithm (Lahr, 1999) and are shown color coded by year. The diamond indicates the volcanic summit location, and black triangles indicate station locations. The catalog locations for the two ML 4.3 events are shown (black diamonds) with their respective focal mechanisms calculated by Moran et al. (2002).
than 1500 events that occurred near Great Sitkin volcano between 1999 and the end of 2005. Routine earthquake locations at AVO for Great Sitkin are calculated using the HYPOELLIPSE algorithm of Lahr (1999) and a 1D velocity model (Fig. 3; Toth and Kisslinger, 1984). The seismic network consists of six short-period stations, five of which are located on the volcano and one of which is located on a nearby island southeast of Great Sitkin (Fig. 1; Table 1). Only station GSTD has three components; all other stations are vertical component only. In addition to the local network, many Great Sitkin events were recorded on three regional stations on or near Adak Island (Fig. 1) and on stations
comprising the network of nearby Kanaga volcano, located ∼75 km west-southwest of Great Sitkin. Catalog locations at Great Sitkin occur in three main areas: (1) ∼10–15 km west and offshore of the volcano, (2) on the southeast flank of the volcano, and (3) in the vicinity of the summit (Fig. 2). Although events have occurred regularly near the summit since 1999 (Fig. 4a), the majority of seismic events at Great Sitkin occurred during a period of increased activity in 2002. Two separate off-summit sequences of events contributed over 750 earthquakes or ∼50% of the total located seismicity over the 6 yr study period. The first sequence occurred in March and April of 2002 (Fig. 4b)
3D P-Wave Velocity Structure and Precise Earthquake Relocation at Great Sitkin Volcano, Alaska
Figure 3.
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Comparison of 1D velocity models used in this study.
and began with an ML 4.3 event on 18 March (Moran et al., 2002). This sequence included over 300 events located 15– 20 km west of the Great Sitkin summit at depths of 10–25 km (Fig. 2). The second and bigger sequence occurred from May through July of 2002 (Fig. 4c). The largest event of this sequence (ML 4.3) occurred on 28 May, several hours after the initial activity. This sequence locates on the southeast volcano flank at depths of 5–15 km (Fig. 2). Almost all events associated with the 2002 sequences are VT earthquakes (Fig. 5). The two ML 4.3 earthquakes are among the largest events ever recorded at an AVO-monitored volcano (Moran et al., 2002). The remaining catalog events have smaller magnitudes (ML < 2:9).
Table 1 Great Sitkin Local and Regional Seismic Station Locations Station
Latitude (deg)
Longitude (deg)
Elevation (m)
ADAG ADK ETKA GSCK GSIG GSMY GSSP GSTD GSTR
51.9802 51.8837 51.8619 52.0119 51.9864 52.0432 52.0928 52.0559 52.0943
176:602 176:685 176:406 176:162 175:925 176:056 176:176 176:145 176:059
286 116 290 384 407 418 295 873 536
Figure 4. Histograms showing the number of events per time interval for (a) the summit region (annual), (b) the first cluster located offshore, west of Great Sitkin beginning March 2002 (daily), and (c) the second cluster beginning May 2002 located southeast of the summit (daily; see Fig. 2 for event locations).
Methodology BS Verification Waveform CC has been successfully used by many researchers to measure the similarity of two time series and to determine precise relative arrival times (e.g., Got et al., 1994; Dodge et al., 1995; Nadeau et al., 1995; Gillard et al., 1996; Shearer, 1997; Rubin et al., 1998; Waldhauser and Ellsworth, 2000; Rowe et al., 2002; Schaff et al., 2004). In order to select only high-quality data, a threshold value Ccut is often chosen (usually 0.7–0.8), below which all data are simply discarded. The choice of Ccut involves a trade-off between limited numbers of very accurate time delays versus a larger number of time delays with some potentially unreliable data included (Du, Thurber, and Eberhart-Phillips, 2004). These traditional CC techniques compare waveforms in the secondorder spectral domain, or power spectrum. If, however, the
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Figure 5.
Seismograms for a typical VT event at Great Sitkin recorded by the local network 29 May 2002. Analyst picks are shown where available. For the three-component station GSTD (bottom), vertical, east, and north components are shown from top to bottom.
waveforms are contaminated by correlated Gaussian or lowskew noise, then traditional methods may provide low CC values for similar events. In this case, comparison in the third-order spectral, or BS domain, where such noise sources are essentially eliminated, may provide more reliable time delays (Du, Thurber, and Eberhart-Phillips, 2004). For waveforms with low signal-to-noise ratios and/or uncorrelated noise, the BS method does not necessarily provide better results than traditional CC techniques. Because neither method always outperforms the other, using the BS method to check the results of traditional CC methods is a valid approach and a way to obtain quality control over the resulting time delays. Such BS verification helps ensure the accuracy of CC results and decreases the ambiguity caused by selecting an appropriate Ccut value. In this approach, two time delays calculated by the BS method, on raw and filtered waveforms, are used to
select or reject the traditional CC time delay. If the three time delays are within a specified range of each other (Δ lim; Table 2), then the CC time delay is successfully verified by the BS method and is included in the relocation process. If the three time delays are not within the specified range (Δ lim), then the time delay is deemed inaccurate and is discarded. This approach ensures that only high-quality time delays are used to relocate hypocenters, increasing the precision of the relocations. We applied the BS verification method to 1535 events that had both waveforms and arrival picks at multiple Great Sitkin stations. Waveforms were band-pass filtered between 1 and 10 Hz, and we adopted time-delay selection criteria values similar to Du, Thurber, and Eberhart-Phillips (2004) (Table 2). These values were chosen by a qualitative visual comparison of waveforms for the Great Sitkin dataset and by
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Table 2 BS Verification Selection Criteria Values for Great Sitkin Criterion
Value
Explanation
Δ lim cc liml cc limu cc lim
1.00 0.40 0.90 0.75
Maximum time delay (sec) difference (CC–BS) allowed Minimum Cmax value considered for verification If Cmax is above this value, all station time delays above cc liml are verified If Cmax is above cc liml but below cc limu, then only stations with CC values above this value are verified
For a more detailed explanation, see Du, Thurber, and Eberhart-Phillips (2004).
testing the results against changes in the individual parameters. We found these values to be relatively robust, such that small changes in the values produce negligible differences in relocation patterns. CCs were performed over a 1 sec data window (50 samples), 0.3 sec before and 0.7 sec after the P pick, yielding 122,660 P-wave time delays. Because of the large uncertainty associated with S-wave picks, a 4.2 sec data window was used for CC, starting 0.9 sec before the S pick. Only ∼13% of the catalog events have catalog S picks, and CC of these yielded 84,878 time delays. S-wave data are important for controlling the depth versus origin time trade-off (Gomberg et al., 1990), and in order to increase the number of S-wave time delays, we used the available S picks and a modified 1D velocity model for Great Sitkin (Fig. 3; dashed line) to estimate appropriate time windows for unpicked S-wave arrivals (e.g., Du, Thurber, Reyners, et al., 2004). This allowed us to obtain differential S times even when there was no catalog S pick. We attempted to pick S on all stations; only station ADK yielded no new S picks. In this way, an additional 12,407 S differential times were added, providing a total of 97,285 S-wave time delays, an increase of ∼15%. Including these differential S times greatly increased event clustering as compared to including differential P times only, especially for those events located on the southeast flank that occurred in 2002. A total of 219,945 time delays selected by BS verification and 469,420 catalog time delays were included in the DD tomography inversion (Fig. 6). Figure 7 shows examples of how the BS verification improves waveform alignment for each of the seismically active areas of Great Sitkin. DD Tomography We first obtained an initial 1D velocity model at Great Sitkin using simul2000 (Thurber and Eberhart-Phillips, 1999; Fig. 3) in order to convert from a layer-based model to a node-based model with linear interpolation between nodes. Using the result as our starting model, we incorporated CC-derived differential times obtained by BS verification, in addition to absolute and differential catalog data, into the DD tomography algorithm tomoDD of Zhang and Thurber (2003) to solve for event locations and 3D velocity structure. A hierarchical weighting scheme was applied in which the catalog data were initially heavily weighted in the inversion to control large-scale relocations (Waldhauser, 2001; Zhang and Thurber, 2003, 2006). Subsequent itera-
Figure 6.
Histograms showing the number of time delays per
CC coefficient selected by the BS verification process at each sta-
tion. All time delays below the chosen threshold value of 0.75 are high-quality time delays that would have been rejected by traditional CC methods (see explanation in text). Note the different vertical scale for GSSP, ADAG, and ETKA. P-wave time delays are shown by gray bars; S-wave time delays are shown by clear bars.
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Figure 7. Examples of BS verification improving waveform alignment at Great Sitkin. For (a)–(d), the upper left panel shows waveforms aligned on preliminary analyst picks, the upper right panel shows waveforms realigned by CC, the lower left panel shows the preliminary waveform stack, and the lower right panel shows the realigned waveform stack. Significant improvement is seen in realigned waveforms and their stacks for (a) southeast cluster events at GSTD, (b) summit events at GSTR, (c) southeast cluster events at GSSP, and (d) west cluster events at GSIG. Similarity is highest at (a) GSTD, where waveform alignment is shown by gray-scaled amplitude intensity due to the greater numbers of similar events there.
3D P-Wave Velocity Structure and Precise Earthquake Relocation at Great Sitkin Volcano, Alaska tions down-weighted catalog data and increased weighting for the differential CC data in order to refine relative event locations. Because of the larger amount and higher precision of P arrival data as compared to S arrival data, here we discuss only the results of the P-wave velocity model. The S-wave velocity model provided small areas with sufficiently high resolution for interpretation, and in these areas, results were in general agreement with those of the P-wave model and provided no new additional information. The 3D P-wave model values from the inversion depend on the grid spacing, the damping, and the smoothing. We placed nodes in our grid so as to achieve uniform node sampling by the ray paths; nodes in areas with a high density of ray paths were more closely spaced than areas with a lower density of ray paths (Fig. 8). To determine the optimal damping and smoothing parameters, we analyzed variance trade-off curves for a range of damping and smoothing values (Eberhart-Phillips, 1986). We chose a damping value of 100 and a smoothing weight of
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15 (in all three spatial directions) as the values that best minimized data variance while maintaining low model variance. Velocity Model Resolution To estimate model resolution, we used both synthetic tests and the derivative weight sum (DWS) of Thurber and Eberhart-Phillips (1999), which is a proxy for ray density. We calculated the average of the nonzero values and 1% of the maximum DWS value in order to estimate resolution from the DWS matrix. Above these thresholds the model resolution is sufficiently high to allow confidence in the inversion results. These indicators are shown by contour lines in model result figures. To test the response of the inversion to velocity perturbations, a synthetic 3D velocity model was used to calculate arrival time data using the same source-receiver distribution as the real data. To assess amplitude recovery and the degree of smearing, we used input anomalies of various geometries and velocities that broadly represented the characteristics of possible heterogeneities in the study
Figure 8. Regional map of Great Sitkin volcano and vicinity showing tomography grid nodes (squares). The positions of the cross sections (Figs. 9 and 10) are shown as distance along the y axis from the grid origin located at the summit (diamond); several x-axis nodes are also labeled for reference in Figures 9 and 10. The inset rectangles show the locations of Figures 11, 12, and 13, which correspond to the seismically active summit, southeast flank, and western offshore regions, respectively. The location of the mapview model slices shown in Ⓔ Figure S1 in the electronic edition of BSSA is indicated by the rotated rectangle enclosing the grid nodes.
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Figure 9. (a) Input velocity model used for synthetic testing of amplitude recovery and smearing. Anomalies are designed to broadly represent characteristics of features in the real model. Alternating fast and slow node () anomalies are also included (Y 4, 5, and 6). Recovery of the synthetic model is shown as (b) absolute V p anomaly relative to the input model (a) and (c) absolute V p (km=sec). DWS contours are also shown: 1% of the maximum DWS value as the outer contour and the mean nonzero DWS value as the inner contour. Position of the cross sections is shown in Figure 8 with the origin at the summit. (Continued)
region, including deep vertical planar faultlike anomalies and a shallow chamberlike feature (Fig. 9a). Constant station noise (0:05 sec) and random arrival time noise (0:02 sec) were added to the synthetic data assuming systematic errors are greater than random ones, and the in-
version was calculated with the same parameterization as the real model. Significant smearing occurs in the synthetic models. This is illustrated in Figure 9b and c, where changes in velocity are 5%–10% at the western end of the model, west of
3D P-Wave Velocity Structure and Precise Earthquake Relocation at Great Sitkin Volcano, Alaska
Figure 9.
any input anomaly. Closer to the volcano, smearing is less significant and generally < 5%. Areas with significant smearing and/or loss of amplitude generally correspond to areas with DWS values below the indicator values, but not exclusively. No input feature is fully recovered. Loss of amplitude of the input anomalies is generally minor, however, but can be as large as 25% (Fig. 9b). Although the synthetic tests (Fig. 9 and Ⓔ Fig. S1 in the electronic edition of BSSA) indicate less-than-ideal resolution for our dataset, this is not surprising given the relatively sparse source-receiver distri-
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Continued.
bution. However, the resolution is sufficient to interpret large-scale heterogeneities with confidence.
Results Tomography There are two significant large-scale velocity anomalies in areas considered well resolved that warrant further discussion (Fig. 10). First, a low-velocity region lies beneath the
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Figure 9.
summit, extending from the surface to ∼10 km depth. This body has appreciable thickness but is not well defined north of the summit where resolution decreases. Low velocities also exist below 10 km, although the geometry of the feature varies somewhat between different cross sections (Fig. 10; Y 0, 1, and 2 km). The second feature is located 10–20 km west of the summit and appears as a region of high velocities abutting a region of low velocities with a vertically
Continued.
oriented trend (near X 18 km). This contrast is visible to some degree in all of the cross sections shown in Figure 10 and primarily occurs in well-recovered regions. Although smearing and imperfect recovery are no doubt affecting this anomaly and prevent a more detailed description of its geometry, the contrast is inferred to be real. Synthetic tests suggest it may be described as a broadly planar, vertically oriented feature (Fig. 9).
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Figure 10.
P-wave velocity model results, earthquake relocations, and DWS contours (discussed in text). Position of the cross sections is shown in Figure 8 with the origin at the summit.
Relocations Figure 2 shows a regional mapview and cross sections of hypocenters located in the vicinity of Great Sitkin. Using only those events with at least four recorded P-wave arrivals and an azimuthal gap in station coverage less than 180°, we were able to relocate 852 hypocenters from the AVO catalog for Great Sitkin. Final root mean square residuals for CC and catalog data are 0.009 and 0.022 sec, respectively. TomoDD
uses the LSQR algorithm (Paige and Saunders, 1982) to solve the damped least-squares problem, which does not readily allow for determination of spatial uncertainties in hypocenter locations. In order to obtain meaningful uncertainty estimates, and as an additional check on the relocations, we used hypoDD on small subsets of events and solved for the relative locations directly using the singular value decomposition (SVD) method, which allows estimation of least-squares uncertainties (Waldhauser and Ellsworth,
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2000; Waldhauser, 2001). Relative spatial uncertainties (2-σ) vary considerably according to region (see the following discussion), but for the entire dataset, they range from less than 10 m to almost 1 km. To confirm these uncertainties, we changed the initial conditions in hypoDD by changing either the number of stations (and hence the station distribution), the initial locations, or the velocity model, and we found location changes comparable to the uncertainties obtained by SVD. In Figures 11, 12, and 13, events using > 40 total differential times (P and S, catalog, and CC) are shown as black circles, and events using < 40 differential times are shown as gray circles.
Figure 11.
Summit Events. Near-surface events are difficult to locate accurately due to the lack of nearby stations and/or inaccurate near-surface velocities (Waldhauser, 2001). This results in large errors in the catalog depths and also can result in events locating above the surface (airquakes). For such events, the AVO algorithm fixes depths 3 km above mean sea level. When relocated using tomoDD, many of these near-surface events became airquakes and were hence discarded. To address this problem, catalog events at < 0 km depth were set to new initial depths below sea level in order to find the depth that resulted in fewer events being lost as airquakes. We tested depths from 1 to 4 km and found the
Initial locations and relocations for summit earthquakes at Great Sitkin. Initial locations are shown only for those events that have been relocated. Squares indicate LP events. The events with initial locations at 1 km depth have been moved from their original catalog locations as discussed in text. Relocations are divided into two quality groups depending on the number of differential times (dts) used. Mapview shows the volcanic summit (diamond), nearby stations (black triangles), and 1 arc-second SRTM topography.
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Figure 12.
Original catalog locations and their relocations for events located on the southeast flank of Great Sitkin, most of which are from the May 2002 swarm. See Figure 11 for additional map details.
resultant seismicity patterns to be generally consistent for different initial depths. We arbitrarily reset events located < 0 km in the catalog to initial depths of 1 km prior to inversion with tomoDD. We retained 212 events that had formerly been airquakes, and the majority of these locate between the summit (1:74 km) and 4 km depth (Fig. 11). For well-located events between 2 and 4 km depth, uncertainties average 21 m horizontally and 35 m vertically. However, uncertainties are likely considerably larger for events at the summit because the highest elevation station is at 0:873 km.
Catalog locations for summit events form a diffuse cloud of seismicity from 2 to 7 km depth centered slightly southeast of the summit (Fig. 11). Our relocations indicate that most of these events are part of a large group located ∼1 km southeast of the summit at ∼3 km depth (Fig. 11). This group of earthquakes shows a rough northeast– southwest linear trend with a shallow northeast plunge. Two other groups of events occur closer to the summit, one just east and the other just west, at depths from 1:74 to > 5 km. The majority of near summit events occurred in late 2000 through early 2001, and the number of events per year steadily declines following this peak in activity (Fig. 4a).
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Figure 13.
Original catalog locations and their relocations for events located west and offshore of Great Sitkin, most of which are part of the March 2002 cluster.
Southeast Events. The majority of the events located beneath the southeast flank of the volcano, and those which lie to the northwest and toward the summit, belong to the event sequence that began 28 May 2002. Relocations indicate clustering not evident in the catalog locations (Fig. 12). Spatial uncertainties average 10 m horizontally and 14 m vertically. The relocations reveal a tight cluster of events along a planar zone oriented northwest–southeast and dipping vertically. This planar zone is located ∼10 km southeast of the summit at depths of 7–11 km. To assess the width of this zone, we performed a least-squares orthogonal distance regression on all the events in close proximity to the large ML 4.3 earthquake that occurred slightly after the sequence began. The best-fit plane for these events has a 2-σ misfit of
426 m. Given the horizontal uncertainties, these events illuminate a broadly planar seismically active region not previously identified at Great Sitkin. Examination of waveforms from this zone reveals S-wave move-out with depth, indicating that the vertical extent of this zone is real and is not an artifact of relocation (Fig. 14). Northwest of this cluster, toward the summit, several other events occurred along the same general azimuth as the trend of the southeast cluster, suggesting a possible connection to the summit. Western, Offshore Events. Station coverage for the sequence of events that began 17 March 2002 off the west coast of Great Sitkin falls far short of ideal (Fig. 1), testing the limits of an already difficult relocation problem. As ex-
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Figure 14.
Sample group of similar waveforms recorded at station GSTD, aligned by P-wave arrival, located southeast of the summit and plotted by increasing relocated depth (shown in kilometers noted in the upper right corner). The separation between the P arrival and the Swave energy increases with depth, indicating that the vertical extent of the relocations is real and not an artifact of the relocation process.
pected, the improvement in hypocenter location for these events is not as great as the improvements seen for the summit and southeast flank events. Spatial uncertainties are noticeably larger than the other two regions, averaging 40 m horizontally and 62 m in depth. Significant improvement is still achieved, however, compared to the original locations (Fig. 13). Although no conclusive spatial or temporal patterns are revealed, the relocations show a significant reduction in the scatter, collapsing a diffuse cloud of seismicity to a much tighter group of events that we speculate lies within a northwest–southeast striking planar zone as discussed in the following section.
Discussion Our relocations of the summit events show that recent seismicity is mostly shallow (< 5 km) and located within a 2 km radius of the summit in mapview (Fig. 11). The small areal extent and wide range of depths for these events (∼7 km) suggest that they may be occurring along an active or previously active volcanic pipelike conduit, which may
connect the summit crater to a deeper magma storage body. The concentration of summit events toward the southeast, at ∼3 km depth and over several years, suggests continued strain accumulation and release that could be related to a continued intrusion or hydrothermal activity. The shallow lowvelocity anomalies and the nature of seismicity at the summit (Fig. 10, Y 0, 1 km depth) suggest that the active hydrothermal system beneath Great Sitkin likely causes these events. Great Sitkin has active fumaroles, mud pots, and hot springs located south of the summit (Fig. 1) and older inactive hydrothermal areas located northeast of these (Motyka et al., 1994). In addition, ∼80 LP events have been recorded at Great Sitkin since installation of the network. Most of these events have catalog locations, although poorly constrained, of less than 5 km in depth within the vicinity of the summit. We are able to relocate only 18 of these LP events, 12 of which occur near the summit (Fig. 11). The location of LP events, the ongoing venting at fumaroles, and the lack of any temporal pattern in the relocations, which have been shown to be precursors to eruptions at other volcanoes (e.g., White, 1996), suggest that current activity at the summit is
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hydrothermal in origin and not due to new or ongoing magmatic activity. The low-velocity regions at > 5 km depth beneath Great Sitkin have P-wave velocities (Fig. 10) slower than global average values for the expected rock types at these depths (Christensen and Mooney, 1995) and for those determined farther east along the Aleutian arc (Holbrook et al., 1999). These areas may represent high temperature and/or partially molten rock in areas of magma storage. In this view, these regions hold the magma that must episodically feed eruptions at Great Sitkin, presumably through the seismically active conduit partially outlined by the summit relocations. Although much less common, LP events have also occurred at much greater depths below Great Sitkin (> 20 km). Power et al. (2004) suggest that these deep LP events probably represent magma from the upper mantle or lower crust feeding shallower magma bodies, which in turn episodically feed the shallower volcanic systems as eruptions or intrusions. This is consistent with our idea that reduced velocities may be indicative of the magma storage system at Great Sitkin. Our data prevent a more detailed investigation of the volcanic plumbing, and our interpretation remains speculative. Seismicity at the summit significantly decreased from 2001 to 2002 and continued to decrease to its lowest level through 2004, before slightly increasing in 2005 (Fig. 4a). The drop in summit seismicity rate began in July 2001, some 8 months prior to the March 2002 and May 2002 sequences. In early 2002, the AVO changed its earthquake detection systems from IASPEI (International Association of Seismology and Physics of the Earth’s Interior) to Earthworm (Dixon et al., 2005), but this change was expected to increase the number of events detected and located, not decrease detection and location. This rate decrease at the summit differs from the more common situation in which flank activity precedes increased summit activity and eruption (e.g., Harlow et al., 1996). The drop in summit seismicity may be purely coincidental, or it may be indicative of a stress change in the volcanic or hydrothermal systems that resulted in a decrease in summit seismicity and an increase in activity beyond the summit region in 2002. Linking seismicity rate changes to stress changes in the volcanic system is speculative but reasonable for the sequence of events located on the southeast flank. The short duration and high spatial correlation of this activity is consistent with a volcanic swarm. The relocations of the sequence suggest a volcanic dike geometry, with a vertical plane oriented radially from the summit and in the direction of regional convergence (320°; Fig. 11). Dike intrusion at depth often trends parallel to the regional maximum compressive stress (Nakamura et al., 1977). Surface geology shows that the locations of several Holocene lava domes northwest of the summit also exhibit this northwest–southeast trend (Waythomas et al., 2003), supporting this orientation as the preferred direction of recent and current magmatic activity. In addition, two episodes of tremor, lasting 20 and 60 min, respectively, occurred ∼20 hr after the initiation
of the swarm. Because tremor is indicative of fluid transport processes, this provides further evidence that these events may be intrusion related (Moran et al., 2002), so we categorize this sequence as a swarm. Furthermore, all of the ∼50 events that occurred prior to the ML 4.3 earthquake relocate at greater depths than later seismicity. These observations fit the model of Roman et al. (2004) for dike emplacement at Iliamna volcano in 1996. In this model, strike-slip events are expected to occur with P axes oriented in the direction of regional maximum compressive stress above the dike tip where the tensile stress increase due to dike propagation is thought to unlock crustal faults above the dike, facilitating slip. Within the dike walls faulting should be perpendicular to the regional maximum compressive stress, consistent with the strike-slip focal mechanism of the ML 4.3 event (Fig. 2). This event and the prior deeper events probably result directly from dike intrusion; the later events of the swarm are probably a combination of an aftershock sequence from the ML 4.3 earthquake and the continued effects of dike emplacement on the surrounding wallrock (Roman, 2005). The b-values for the entire swarm (1.1) and for only those events following the ML 4.3 event (1.1) are typical of tectonic aftershock sequences (Fig. 15), but low b-values have also been reported for swarms of VT events at other volcanoes (e.g., Mori et al., 1996; Roman et al., 2004). Without additional supporting data, such as focal mechanisms, it is difficult to attribute this swarm conclusively to an intrusion. However, we believe that enough circumstantial evidence exists to suggest that a dike intrusion on the southeast flank of Great Sitkin volcano likely occurred in 2002. The relocations of the events from the western cluster are not as precise as those of the southeast swarm discussed previously. However, the available data suggest that these events are not magmatic in origin but constitute a mainshock–aftershock sequence on a previously unknown fault. The remoteness of these earthquakes with respect to the nearby active volcanic centers (Fig. 2) and the low b-value (Fig. 15) support this interpretation. However, there is some ambiguity as to the orientation of this fault. The northeast–southwest nodal plane for this sequence’s mainshock focal mechanism (Fig. 2) could be interpreted as dextral strike-slip motion associated with backarc processes caused by the clockwise block rotation of the Aleutian arc massif (Geist et al., 1988). This type of motion along and behind the volcanic arc has previously been observed in the vicinity of Great Sitkin following the 1986 Mw 8.0 Andreanof Islands subduction event and is similarly attributed to such motion (Ekstrom and Engdahl, 1989). Alternatively, and we believe more likely, this fault shows sinistral motion on the northwest–southeast oriented nodal plane, the direction parallel to regional convergence. As an additional test to help determine if these events should in fact locate with a northwest–southeast trend, we measured the S minus P times for these events at station GSTD and plotted them versus distance along a northwest–southeast line and, for comparison, along a northeast–southwest line. Although neither plot
3D P-Wave Velocity Structure and Precise Earthquake Relocation at Great Sitkin Volcano, Alaska
Figure 15. b-values calculated for seismic activity occurring in 2002. (a) All events on the southeast flank, (b) events following the ML 4.3 mainshock on the southeast flank, and (c) events occurring offshore, west of Great Sitkin. Regression lines are plotted from upper to lower magnitude cutoff. A line is also shown for the magnitude of completeness (0.6; Dixon et al., 2005).
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subduction. Although not strictly within the idealized boundaries of the Andreanof block defined in their model, it is possible that this fault west of Great Sitkin is a developing fault accommodating intrablock deformation within this largest of the Aleutian arc blocks. Geist et al. (1988) and Ryan and Scholl (1993) note that the western Andreanof block currently appears unrotated, leaving unanswered the question of how the oblique component of subduction is being accommodated in this region. We speculate that this fault may be an extensive crustal feature developing in response to rotation of the western part of the Andreanof block and that it may be responsible for the sag basins and the channel cut seen in Figure 16 along the same trend we infer from the relocations. If so, both normal and dextral strike-slip faulting may be expected, consistent with the mainshock focal mechanism for the March 2002 event. Our P-wave velocity model also supports the presence of a fault west of Great Sitkin, running parallel to convergence, and that this fault may be pervasive throughout the study region. Figure 10 shows that there is a distinct, vertically oriented velocity contrast located in the same region as the March 2002 event cluster. Our synthetic tests (Fig. 9) indicate that a narrow vertical low-velocity zone could be successfully imaged in that area. Although smearing is considerable, it is clear that such a velocity contrast could be resolved in our model, and therefore the velocity contrast we see is likely a real feature that may be interpreted as a near vertical fault. The exact geometry of the fault is unclear, but the contrast exists to some degree throughout most of the model, oriented in the direction parallel to convergence. The model resolution wanes near the edges, but it is possible that this feature is pervasive and extends beyond the boundaries of the study area in the same orientation. Alternatively, this area may represent a change in lithology. If an area of weakness exists between these opposing rock types, then faulting may preferentially occur here. Unfortunately, other geophysical data do not exist in high enough resolution to confirm the existence of such a fault zone, and our interpretations remain speculative.
Conclusions shows a good linear trend, we find that the northwest– southeast line provides a slightly better fit to the S minus P times, which supports our interpretation of these events as more likely occurring along a northwest–southeast oriented fault. Detailed bathymetry of the region provides additional evidence that a fault oriented parallel to the convergence direction may exist in this area (Nichols and Perry, 1966). Figure 16 shows that the relocations of these events coincide with three small basins, two of which are elongated in the convergence direction. The Aleutian arc block-rotation model of Geist et al. (1988) describes how block-bounding tear canyons develop throughout the arc as a result of oblique
Small seismic networks such as the one at Great Sitkin volcano provide difficult conditions in which to obtain reliable locations. Despite such difficulties, our combined BS verification and DD tomography method significantly improves the view of seismicity at Great Sitkin and provides greater insight into its active volcanic system. Relocations of catalog data reveal a hydrothermally active summit conduit, a possible dike intrusion, and an active offshore crustal fault. P-wave velocity modeling images low-velocity regions directly beneath the summit of Great Sitkin and at depth, which may be interpreted as regions containing hydrothermal fluids and areas of magma storage, respectively. West and offshore of Great Sitkin, a zone of anomalous velocities
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Figure 16. Bathymetric map of the Andreanof Islands from Nichols and Perry (1966), contour interval is 50 fathoms (∼90 m). Overlain on the map are the relocations of this study for the March 2002 sequence and its mainshock focal mechanism determined by Moran et al. (2002), the idealized boundary of the Andreanof block (dashed line: Geist et al., 1988), and the location of the 1957 Mw 8.6 earthquake. The bathymetry shows three small basins and a channel cut that lie along a trend line (dotted line) approximately parallel to the convergence direction (arrow) and our inferred trend for the events of the sequence. The inset box shows the Great Sitkin region bathymetry in greater detail.
3D P-Wave Velocity Structure and Precise Earthquake Relocation at Great Sitkin Volcano, Alaska is located in the same region as the earthquake cluster of March–April 2002. This zone may represent an extensive vertical crustal fault oriented parallel to convergence. Our results provide another example showing that highprecision earthquake locations can be obtained in volcanic regions despite sparse station coverage and less-than-ideal signals (e.g., Jones et al., 2001; Battaglia et al., 2004; Rowe et al., 2004). The additional quality check on the CC-derived differential times provided by BS verification decreases the chances of producing unreliable locations and/or artifacts and suggests that such a method may be applicable to volcanic hazard monitoring in real or near real time. Our relocations and velocity modeling at Great Sitkin and complementary studies at other Alaska volcanoes (Brown et al., 2006; DeShon et al., 2007) demonstrate the ability of these methods to delineate seismic features and trends at active volcanoes reliably, which can provide important information for future volcano monitoring.
Data and Resources The Great Sitkin network catalog data were provided by the Alaska Volcano Observatory (AVO) and are available upon request (contact James Dixon,
[email protected]). Many of the figures were produced using the Generic Mapping Tools software (www.soest.hawaii.edu/gmt; Wessel and Smith, 1998). In addition, data from the Shuttle Radar Topography Mission (SRTM) was used in two figures (data are freely available at http://srtm.usgs.gov/, last accessed December 2006); however, this was not a main data source for the study.
Acknowledgments We thank James Dixon and the Alaska Volcano Observatory (AVO) for access to the Great Sitkin network catalog data, and we thank Seth Moran and John Power for additional data and discussions. This manuscript benefited greatly from the thorough review of an anonymous reviewer and from comments by Guoqing Lin, Brad Singer, and Brian Jicha. This material is based upon work supported by the National Science Foundation under Grant Number EAR-0409291, awarded to C. T.
References Battaglia, J., C. H. Thurber, J.-L. Got, C. A. Rowe, and R. A. White (2004). Precise relocation of earthquakes following the 15 June 1991 eruption of Mount Pinatubo (Philippines), J. Geophys. Res. 109, B07302, doi 10.1029/2003JB002959. Brown, J., H. Deshon, S. Prejean, C. Thurber, and J. Power (2006). Crosscorrelation analysis reveals waveform similarity in long-period events prior to eruptive activity at Mt. Spurr volcano, Alaska (abstract), Seism. Res. Lett. 77, 241. Christensen, N. I., and W. D. Mooney (1995). Seismic velocity structure and composition of the continental crust: a global view, J. Geophys. Res. 100, 9761–9788. DeShon, H. R., C. H. Thurber, and C. Rowe (2007). High-precision earthquake location and three-dimensional P wave velocity determination at Redoubt volcano, Alaska, J. Geophys. Res. 112, B07312, doi 10.1029/ 2006JB004751.
2447
Dixon, J. P., S. D. Stihler, and J. A. Power (2005). A comparison of seismic event detection with IASPEI and Earthworm acquisition systems at Alaskan volcanoes, Seism. Res. Lett. 76, 168–176. Dodge, D. A., G. C. Beroza, and W. L. Ellsworth (1995). Foreshock sequence of the 1992 Landers, California, earthquake and its implications for earthquake nucleation, J. Geophys. Res. 100, 9865–9880. Du, W.-X., C. H. Thurber, and D. Eberhart-Phillips (2004). Earthquake relocation using cross-correlation time delay estimates verified with the bispectrum method, Bull. Seismol. Soc. Am. 94, 856–866. Du, W.-X., C. H. Thurber, M. Reyners, D. Eberhart-Phillips, and H. Zhang (2004). New constraints on seismicity in the Wellington region of New Zealand from relocated earthquake hypocentres, Geophys. J. Int. 158, 1088–1102. Eberhart-Phillips, D. (1986). Three-dimensional velocity structure in northern California coast ranges from inversion of local earthquake arrival times, Bull. Seismol. Soc. Am. 76, 1025–1052. Ekstrom, G., and E. R. Engdahl (1989). Earthquake source parameters and stress distribution in the Adak Island region of the Central Aleutian Islands, J. Geophys. Res. 94, 15,499–15,519. Geist, E. L., J. R. Childs, and D. W. Scholl (1988). The origin of summit basins of the Aleutian ridge: implications for block rotation of an arc massif, Tectonics 7, 327–341. Gillard, D., A. M. Rubin, and P. Okubo (1996). Highly concentrated seismicity caused by deformation of Kilauea’s deep magma system, Nature 384, 343–346. Gomberg, J. S., K. M. Shedlock, and S. W. Roecker (1990). The effect of S-wave arrival times on the accuracy of hypocenter estimation, Bull. Seismol. Soc. Am. 80, 1605–1628. Got, J. L., J. Frechert, and F. W. Klein (1994). Deep fault plane geometry inferred from multiplet relative relocation beneath the south flank of Kilauea, J. Geophys. Res. 99, 15,375–15,386. Harlow, D. H., J. A. Power, E. Laguerta, G. Ambubuyog, R. A. White, and R. P. Hoblitt (1996). Precursory seismicity and forecasting of the June 15, 1991 eruption of Mount Pinatubo, Philippines, in Fire and Mud: Eruptions and Lahars of Mount Pinatubo, Philippines, C. G. Newhall and R. S. Punongbayan (Editors), University of Washington Press, Seattle, 285–305. Holbrook, W. S., D. Lizarralde, S. McGeary, N. Bangs, and J. Diebold (1999). Structure and composition of the Aleutian island arc and implications for continental crustal growth, Geology 27, 31–34. Jones, J. P., C. H. Thurber, and W. J. Lutter (2001). High-precision location of pre-eruption seismicity at Mount Pinatubo, Phillipines, 30 May– 3 June, 1991, Phys. Earth Planet. Inter. 123, 221–232. Lahr, J. C. (1999). HYPOELLIPSE: A computer program for determining local earthquake hypocentral parameters, magnitude, and first motion pattern, U.S. Geol. Surv. Open-File Rept. 99-23, 116 pp. McNutt, S. R. (2005). Volcanic seismology, Annu. Rev. Earth Planet. Sci. 33, 461–491. Miller, T. P., R. G. McGimsey, D. H. Richter, J. R. Riehle, C. J. Nye, M. E. Yount, and J. A. Dumoulin (1998). Catalog of the historically active volcanoes of Alaska, U.S. Geol. Surv. Open-File Rept. 98-0582, 104 pp. Moran, S. C., J. A. Power, T. A. Plucinski, J. F. Paskievitch, and J. P. Dixon (2002). Seismicity at Great Sitkin volcano, Andreanof Islands, Alaska (Abstract S11C-1163), EOS Trans. AGU 83 (Fall Meet. Suppl.), S11C-1163. Mori, J., R. A. White, D. H. Harlow, P. Okubo, J. A. Power, R. P. Hoblitt, E. P. Laguerta, A. Lanzuza, and B. C. Bautista (1996). Volcanic earthquakes following the 1991 climactic eruption of Mount Pinatubo: strong seismicity during a waning eruption, in Fire and Mud: Eruptions and Lahars of Mount Pinatubo, Philippines, C. G. Newhall and R. S. Punongbayan (Editors), University of Washington Press, Seattle, 339–350. Motyka, R. J., S. A. Liss, C. J. Nye, and M. A. Moorman (1994). Geothermal resources of the Aleutian arc, Alaska Division of Geological and Geophysical Surveys Profess. Rept. 114, 17 pp., 4 sheets, scale 1:1,000,000.
2448
J. D. Pesicek, C. H. Thurber, H. R. DeShon, S. G. Prejean, and H. Zhang
Nadeau, R. M., W. Foxall, and T. V. McEvilly (1995). Clustering and periodic recurrence of microearthquakes on the San Andreas fault at Parkfield, California, Science 267, 503–507. Nakamura, K., K. H. Jacob, and J. N. Davies (1977). Volcanoes as possible indicators of tectonic stress orientation: Aleutians and Alaska, Pure Appl. Geophys. 115, 87–112. Nichols, H., and R. B. Perry (1966). Bathymetry of the Aleutian Arc, Alaska, U.S. Coast and Geod. Surv. Monograph 3, scale 1:400,000 Paige, C. C., and M. A. Saunders (1982). LSQR: sparse linear equations and least squares problems, ACM Trans. Math. Softw. 8, no. 2, 195–209. Power, J. A., S. D. Stihler, R. A. White, and S. C. Moran (2004). Observations of deep long-period (DLP) seismic events beneath Aleutian Arc volcanoes: 1989–2002, J. Volcanol. Geotherm. Res. 138, 243–266. Roman, D. C. (2005). Numerical models of volcanotectonic earthquake triggering on non-ideally oriented faults, Geophys. Res. Lett. 32, L02304, doi 10.1029/2004GL021549. Roman, D. C., J. A. Power, S. C. Moran, K. V. Cashman, M. P. Doukas, C. A. Neal, and T. M. Gerlach (2004). Evidence for dike emplacement beneath Iliamna volcano, Alaska in 1996, J. Volcanol. Geotherm. Res. 130, 265–284. Rowe, C. A., R. C. Aster, B. Borchers, and C. J. Young (2002). An automatic, adaptive algorithm for refining phase picks in large seismic data sets, Bull. Seismol. Soc. Am. 92, 1660–1674. Rowe, C. A., C. H. Thurber, and R. A. White (2004). Dome growth behavior at Soufriere Hills volcano, Montserrat, revealed by relocation of volcanic event swarms, 1995–1996, J. Volcanol. Geotherm. Res. 134, 199–221. Rubin, A. M., D. Gillard, and J. Got (1998). A reinterpretation of seismicity associated with the January 1983 dike intrusion at Kilauea volcano, Hawaii, J. Geophys. Res. 103, 10,003–10,015. Ryan, H. F., and D. W. Scholl (1993). Geologic implications of great interplate earthquakes along the Aleutian arc, J. Geophys. Res. 98, 22,135–22,146. Schaff, D. P., G. H. R. Bokelmann, W. L. Ellsworth, E. Zanzerkia, F. Waldhauser, and G. C. Beroza (2004). Optimizing correlation techniques for improved earthquake location, Bull. Seismol. Soc. Am. 94, 705–721. Shearer, P. M. (1997). Improving local earthquake locations using the L1 norm and waveform cross correlation: application to the Whittier Narrows, California, aftershock sequence, J. Geophys. Res. 102, 8269–8283. Thurber, C., and D. Eberhart-Phillips (1999). Local earthquake tomography with flexible gridding, Comput. Geosci. 25, 809–818.
Toth, T., and C. Kisslinger (1984). Revised focal depths and velocity model for local earthquakes in the Adak seismic zone, Bull. Seismol. Soc. Am. 74, 1349–1360. Waldhauser, F. (2001). HypoDD: a computer program to compute double-difference hypocenter locations, U.S. Geol. Surv. Open-File Rept. 01-113, 25 pp. Waldhauser, F., and W. L. Ellsworth (2000). A double-difference earthquake location algorithm: method and application to the Northern Hayward fault, California, Bull. Seismol. Soc. Am. 90, 1353–1368. Waythomas, C. F., T. P. Miller, and C. J. Nye (2003). Preliminary geologic map of Great Sitkin volcano, Alaska, U.S. Geol. Surv. Open-File Rept. OF 03-0036, 1 sheet, scale 1:63,360. Wessel, P., and W. H. F. Smith (1998). New, improved version of Generic Mapping Tools released, Eos Trans. AGU 79, 579. White, R. A. (1996). Precursory deep long-period earthquakes at Mount Pinatubo: spatio-temporal link to a basaltic trigger, in Fire and Mud: Eruptions and Lahars of Mount Pinatubo, Philippines, C. G. Newhall and R. S. Punongbayan (Editors), University of Washington Press, Seattle, 307–327. Zhang, H., and C. H. Thurber (2003). Double-difference tomography: the method and its application to the Hayward fault, California, Bull. Seismol. Soc. Am. 93, 1875–1889. Zhang, H., and C. Thurber (2006). Development and Applications of Double-difference Seismic Tomography, Pure Appl. Geophys. 163, 373–403, Table 1.
Department of Geology and Geophysics University of Wisconsin-Madison 1215 West Dayton Street Madison, Wisconsin 53706 (J.D.P., C.H.T., H.R.D., H.Z.)
U.S. Geological Survey Alaska Volcano Observatory 4210 University Avenue Anchorage, Alaska 99508 (S.G.P.)
Manuscript received 21 August 2007
