On the relation between infrasound, seismicity, and small pyroclastic explosions at Karymsky Volcano

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1 JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 112,, doi: /2006jb004654, 2007 On the relation between infrasound, seismicity, and small pyroclastic explosions at Karymsky Volcano Jeffrey B. Johnson 1 Received 25 July 2006; revised 26 February 2007; accepted 7 May 2007; published 11 August [1] Ground- and atmospheric-propagating elastic waves are integrated with digitized video to reconstruct important parameters during the onset of eruptions at Karymsky Volcano in Muzzle velocities, which range from 10 to 55 m/s for 34 analyzed explosive events, are found to be very well correlated with eruption velocities synthesized from infrasonic trace data. Acoustic modeling proposed here invokes a linear acoustic source consistent with a 3-m piston that accelerates upwards. Timing of acoustic arrivals relative to video records of material emission is used to constrain fragmentation sources depths, which are assumed to be coincident with infrasound generation. These sources are generally shallow, with an observed vertical range of about 15 m. Systematic changes in source depth are observed during two nights of data collection and are attributed to variations in the depth of the magma free surface. Relative timing of acoustic and seismic arrivals is used to identify the presence of seismicity prior to fragmentation. This seismicity is emergent and low intensity, but significant because it precedes the presumed fragmentation events by variable amounts up to 4 s. The primary short-period seismic phases appear coincident with the initiation of fragmentation and infrasound generation. Because of the complexity of the seismic Green s function, it is not possible to correlate amplitude from a specific seismic pulse with observed eruption intensity. Citation: Johnson, J. B. (2007), On the relation between infrasound, seismicity, and small pyroclastic explosions at Karymsky Volcano, J. Geophys. Res., 112,, doi: /2006jb Introduction [2] Erupting volcanoes can radiate substantial elastic energy into the atmosphere and into the ground. Volcano seismoacoustic studies, in which both ground- and atmospheric-propagating waves are recorded, are becoming increasingly common as researchers recognize the importance of monitoring the entire elastic wavefield [e.g., Garces and McNutt, 1997; Hagerty et al., 2000; Rowe et al., 2000; Ripepe et al., 2002; Caplan-Auerbach and McNutt, 2003; Johnson et al., 2004; Ruiz et al., 2006; Moran et al., 2004]. Although seismic radiation provides the best tool for probing motions internal to a volcano, acoustic radiation, with its much simplified Green s function, offers a more direct means to recover physical source motions in the vicinity of the volcanic vent [e.g., Wouff and McGetchin, 1975]. [3] Integrated seismoacoustic studies from the early 1980s at Erebus Volcano demonstrated that volcanic explosions simultaneously emit infrasound and short-period seismicity [Dibble, 1989]. Simple short-duration eruptions, from Karymsky for instance, show a very simple acoustic pressure pulse and corresponding more complex seismic 1 Department of Earth and Environmental Science, New Mexico Institute of Mining and Technology, Socorro, New Mexico, USA. Copyright 2007 by the American Geophysical Union /07/2006JB004654$09.00 signal which arrives prior to the acoustic transient because of higher intrinsic compressional wave velocities in the ground [Johnson et al., 1998]. Other studies have shown that many complex seismic signals, including some earthquakes identified as long-period events and tremor, of both the harmonic and broadband varieties, are also manifested in the acoustic wavefield [e.g., Garces and McNutt, 1997; Hagerty et al., 2000; Johnson and Lees, 2000; Rowe et al., 2000; Caplan-Auerbach and McNutt, 2003]. It has also been shown that long-duration seismic wave trains, with durations greater than the volcano s seismic impulse response, reflect extended-duration degassing source processes [e.g., Johnson and Lees, 2000; Moran et al., 2004]. [4] Current models to explain infrasound generation vary widely. Though several types of physical source mechanisms are theoretically capable of producing identical infrasonic transients by perturbing the atmosphere, this paper favors a model in which explosive expansion of volcanic volatiles is responsible [e.g., Dibble, 1989; Firstov and Kravchenko, 1996; Yamasato, 1998; Ripepe et al., 2001]. At Erebus Volcano, for instance, integrated video and infrasound records have effectively documented that acoustic pulses are coincident with large bubbles ruptured at the surface of a lava lake [Dibble, 1989; Rowe et al., 2000; examples at [5] Elsewhere, other mechanisms have been invoked to explain volcanic infrasound, including rapidly oscillating 1of21

2 Figure 1. (a h) Sequence of video stills of a typical Karymsky explosion (onset of an eruptive event) taken from helicopter. Image time is referenced to T I, the presumed source time of simultaneous magma fragmentation and infrasound generation (video footage courtesy of J. Lees, 1998). large (>10 m) bubbles at the surface of a fluid magma [Vergniolle et al., 2004] or accelerations of the magma free surface induced by immersed sources and/or resonant oscillations of a magma-filled conduit [e.g., Garces and McNutt, 1997]. However, very high amplitude infrasound, which commonly exceeds a few pascals more than 1 km from the vent [Johnson et al., 2004], requires enormous volumetric accelerations, which are more effectively accomplished by explosive gas expansion than by large displacements of a fluid magma free surface. 2. Background [6] Episodic explosive degassing, occurring with frequencies ranging from minutes to hours, represents one of the most common types of eruptive behaviors occurring at andesitic volcanoes worldwide [Simkin and Siebert, 1995]. Some contemporaneous examples of such activity (as of December 2006) include Karymsky (Russia), Volcan de Colima (Mexico), Fuego (Guatemala), Reventador, Sangay, and Tungurahua (Ecuador), and others [Bulletin of the Global Volcanism Network (BGVN ), 2006]. Typically many of these eruptive events begin impulsively and taper over the course of seconds to minutes. Eruption products include gas plumes with variable content of ash and ballistics, which may be composed of both juvenile and recycled material from previous events. At Karymsky Volcano, these frequent events, occurring 10 or more times each hour during 1999, likely represent repeatable magma foam disruption events, which do not individually appear to significantly alter vent/conduit morphology [Johnson et al., 1998; Johnson and Lees, 2000; Fischer et al., 2002; Ozerov et al., 2003]. [7] Characterization of Karymsky-type eruptions in the summer of 1999 is semantically challenging because they do not clearly fall into either the Strombolian or Vulcanian terminology. According to traditional definitions, Strombolian suggests relatively low-vigor, quasi-continuous degassing associated with emission of incandescent ballistics while Vulcanian eruptions imply short blast-like eruptions of gas and ash. According to the Volcanic Explosivity Index (VEI) [Newhall and Self, 1982], Vulcanian eruptions are significantly more intense than Strombolian-type events producing eruptions columns ranging from 1 to 15 km. However, many of the short blast-like eruptions of gas and ash at Karymsky generate ash columns rising no higher than 1 km and would be quantified individually as VEI 0. As such, discrete events at Karymsky are best characterized as small, ballistic and ash-laden eruptive events. Accordingly, Karymsky events have some traits consistent with common definitions of both Strombolian and Vulcanian activity. Eruptive events may be composed of one or more explosive pulses, where explosive pulse is defined here as a shortduration, impulsive outflow. Eruptive events may also include more extended intervals of jet-like emissions or periodic sequences of harmonic tremor pulses, chugging, which may reflect rhythmic gas exhalations [Johnson et al., 2000] or alternatively deflections of the magma free surface [Garces and McNutt, 1997]. [8] During September 1999, Karymsky exhibited frequent eruptive events similar to the activity witnessed during 1998 and 1997 [Johnson et al., 2003]. Unlike the observed 1998 activity, no lava flows were extruded. However, incandescent ballistics and ash attested to a continued, but potentially reduced, flux of juvenile material out through the vent. During a weeklong study interval in September 1999, eruptive events occurred on average eight times each hour. Explosive event duration, as approximated from the length of seismic coda remaining above background, ranged from 20 s to more than a minute. At night, during clear weather, some incandescence was observed for each event. During the day, ashy columns were visible, reaching elevations of up to about 1 km above the vent. Nearly all these events began suddenly, without discrete precursory seismic transients and without precursory (visible) gas emissions. Aerial footage of the Karymsky vent from a typical eruption in 1998 illustrates the sudden emission of gas and pyroclasts from a lava/tephra-choked vent (Figure 1). This rapid expansion of gas is hypothesized 2of21

3 Figure 2. (a) Map and (b) profile of Karymsky stations. (c) Photo of Karymsky from video acquisition site showing approximate field of view (FOV) of digital images. to be responsible for the impulsive, high-amplitude N- shaped infrasonic wavelet common at the onset of each acoustic pressure transient. 3. Data [9] Karymsky eruptive events were recorded with a 4- station infrasonic/seismic network in September Each stand-alone station recorded triaxial broadband seismic and one to three acoustic channels (dependent upon site) to a 6- channel Reftek A07 data acquisition system provided by the Program for Array Seismic Studies of the Continental Lithosphere (PASSCAL). Recording was continuous at 125 Hz. Trace data presented here come exclusively from station KRM3, which was located an estimated 1800 ± 50 m slant distance from the vent (Figure 2). KRM3 was operational from 5 to 12 September [10] Constraints on elastic wave radiation were accomplished through simultaneous acquisition of nighttime video, which provided high-quality video footage of the vent region. All videos were taken from a camera site located on the caldera rim 2300-m slant distance from the vent 3of21

4 Figure 3. Normalized, unfiltered vertical seismic and infrasonic data from the 34 Karymsky explosive events recorded at station KRM3. Listed events 3xx were recorded on the night of September, and events 4xx were recorded on the night of September. and approximately 500 m below the summit (Figure 2). Digital video was acquired on the nights of (20:27 to 04:45 local time) and September (19:37 to 01:55 local time). During these two nights 90 high-quality eruptive events were identified, windowed, and digitized at 0.1-s intervals from the analog video footage. A subset of 34 of these events have been selected for analysis in this study because corresponding infrasound transients are clearly identifiable and possess initial pulses that exceed background noise levels by at least a factor of two. These 34 events recorded at KRM3 are shown in Figures 3 and 4 and are organized chronologically. Zero time is referenced to the onset of the initial infrasonic pressure pulse. Threedigit event numbers are assigned sequentially corresponding to all eruption earthquakes as identified by seismic data Seismic [11] Broadband seismic recordings were made with a CMG40-T three-component broadband seismometer with inband sensitivity of 800 V/m/s and corner frequency at 30 s. For analytical purposes, data were filtered above 1 Hz with a two-pole Butterworth filter to avoid 0.3 Hz microseisms from ocean waves (Figure 5). Above this bandwidth, the sensor response is considered flat. The seismic event highlighted in Figure 5 has a particularly high level of microseism noise relative to signal associated with the 4of21

5 Figure 4. Windowed 12-s version of Figure 3. Vertical velocity and pressure values correspond to peak-to-peak amplitudes as recorded at KRM3. eruptive event and demonstrates the need for high-pass filtering Infrasound [12] Infrasonic recordings were made with a Larsen- Davis 2570 electret condenser microphone with laboratorycalibrated single-pole corner frequency at 0.27 Hz (3 db down) and nominal sensitivity of 48 mv/pa. This sensor is suitable for recording the majority of infrasound radiated by most volcanoes, which generally occupies the near-infrasound bandwidth (1 20 Hz) [Johnson et al., 2004]. The bulk of Karymsky infrasound, which peaks at 1.5 Hz for the data presented here, may be considered well within the sensitivity bandwidth for the Larsen-Davis microphones (Figure 6). Nevertheless, because accurate pressure time series are vital for proper acoustic waveform modeling, an effort is made to deconvolve the Larsen-Davis instrument response, as calibrated inhouse. Digital deconvolution is performed by applying a transfer function to the 125-Hz data with a single zero at , single pole at 1, and gain of In addition, a two-pole 0.05 Hz high-pass filter is applied so that very low frequency drift associated with presumed electronic noise does not overwhelm the deconvolved signal. Figure 6 shows that instrument response deconvolution does not significantly alter the general shape of the primary infrasound pulse, which consists of frequencies primarily 0.5 Hz and higher. It should be noted that the event depicted in Figure 6 has one of the lowest frequency contents of the 34 events showcased in this paper. This event is relatively unique in that it does not have a 5of21

6 Figure 5. (a) Raw (thin line) and filtered (thick line) vertical seismic trace from event #440 recorded at station KRM3. (b) Corresponding frequency spectra for event #440 before and after filtering. (c) Stacked frequency spectra for vertical channel seismicity for 34 eruptive events recorded at KRM3. Amplitude axes for spectral plots are linear. pronounced negative rarefaction following the initial compression Video [13] A Sony CCD-TRU75 Hi8 camera recorded nearly continuous video during two consecutive nights. Viewing conditions were generally excellent in terms of cloud cover, yielding more than 10 hours of footage. Of this footage, approximately 90 eruptive events were identified and windowed. Of these, a 34-event subset was not significantly obscured by fume and had accompanying low-noise infrasound suitable for analysis. The angle of inclination from camera site to the target was 15 above horizontal with a field of view (FOV) encompassing an area of m at the distance of the vent. This perspective implies that vertically rising material is foreshortened by only about 3.5%. All video analysis was carried out on footage shot using the Sony Nightvision mode, which allowed a fixed exposure for all events. For each event, 30 or more seconds of windowed video was digitized ( pixels) at 10 Hz for image processing. Accurate video timing (0 to +0.1 s relative to time base) was achieved by synchronization of the camera s internal clock with a Global Positioning Figure 6. (a) Raw (thin line) and filtered (thick line) infrasonic trace from event #346 recorded at station KRM3. (b) Corresponding frequency spectra for event #346 before and after sensor response deconvolution. (c) Stacked frequency spectra for infrasound from 34 eruptive events recorded at KRM3. Amplitude axes for spectral plots are linear. 6of21

7 Figure 7. (a f) Example of still video image frames (shown at 0.2-s intervals). Time displayed in the video frames is GMT, which is 11 hours behind Kamchatka local time. (g) Corresponding time of synchronized acoustic and (h) seismic waveforms recorded at KRM3. The eruption is observed visually more than 5 s before infrasound reaches station KRM3. System (GPS) handheld unit. An example of seismoacoustic-video synced data is shown in Figure Interpretation [14] The primary motivation of this paper is better understanding of the relation between elastic energy radiation and eruption characteristics for small, frequent, pyroclastladen explosions. Toward this end, four fundamental topics are now investigated in the following order: (1) an estimation of depth and variability in depth of fragmentation or explosion source, (2) a correlation of infrasound-derived eruption velocities with muzzle velocity, (3) an evaluation of the temporal coincidence of seismic and acoustic radiation, and (4) a study of the poor correlation of seismic intensity with eruptive events. [15] To address items 1, 2, and 4, eruption parameters are quantified from the digitized nighttime video and compared with elastic energy radiated into the seismic and acoustic wavefields. The most reliable measurements of eruptive intensity, as extracted from the video, are calculated at the explosive onset of each eruptive event because, within seconds after the onset of each event, an ash or vapor cloud cools and provides an opaque cover to continuing incandescent emissions. As such, visible incandescence from subsequent explosive pulses within the same eruptive event is attenuated. Generally, the vent is clear after a few minutes, in advance of the next episodic eruptive event. The first few seconds of each eruptive event may thus be used to extract important eruption parameters. The most important of these parameters are summarized in Table 1 and discussed in sections 4.1, 4.2, 4.3, and 4.4 for the 34 different events Depth to Fragmentation Within the Conduit [16] The first video frame in which incandescence appears in the FOV of the camera is identified as time of video (T V ) (Figure 8c). Clearly, for an explosion source occurring at some unknown depth within a crater or conduit, a finite time has already elapsed since initial disruption of the magma. This disruption, or fragmentation, is postulated here to coincide with the sudden failure of a highly porous vesicle-rich magma during which time-pressurized volatiles are explosively released to the atmosphere [Sparks, 1998]. This initial fragmentation time (T I ) is shown in this section to be temporally conjoint with the generation of the highintensity, impulsive infrasound compression that characterizes the onset of nearly all Karymsky eruptive events 7of21

8 Table 1. Summary of Eruption Parameters From 34 Analyzed Eruptive Events GMT, hh:mm:ss Event # DT A V, s V c(0.1 s), m/s V c(0.4 s), m/s V c(1 s), m/s DT V I, s DH, m U max, m/s a, m DT P I, s t, s A(t) max, Pa 11 September :27: :53: :00: :12: :27: :24: :38: :24: :37: :51: :57: :10: :08: :32: :49: :29: :35: :57: :01: :22: :34: :44: September :37: :02: :31: :56: :53: :37: :53: :09: :48: :27: :35: :54: S(t) max, mm/s (Figure 8b). In section 4.2, additional evidence is provided to show the coincidence of fragmentation and infrasound generation by demonstrating that infrasound waveform amplitudes are well correlated with the intensity of the eruptive event onset. [17] The conduit transit time DT V I (T V T I ) is defined here as the time required for pyroclasts to travel the length of an open conduit and hidden portion of the crater (i.e., from fragmentation source region to the FOV of the camera). This time is dependent upon the conduit transit distance (DH) and the integrated slowness (1/V c (h)) of the rising material within the conduit and crater at depths h below the camera FOV: DT V I ¼ Z DH 0 dh V c ðþ h The transit time DT V I is not directly measured but can be recovered from the measured value DT A V, which is the time delay between acoustic arrival at KRM3 and the visual onset of incandescence. Combined with eruption muzzle velocity, as estimated from the video, fragmentation onset times may be determined and used to estimate depth in the conduit to the fragmentation/infrasound source. [18] The extraction of muzzle velocity from photographic media has precedent in earlier time-lapse photographic studies for similarly small episodic eruptions. Wilson and ð1þ Self [1980] measured cloud expansion velocities for small Vulcanian eruptions at Fuego Volcano, Guatemala, while both Chouet et al. [1974] and Ripepe et al. [1993] estimated muzzle velocities and kinetic energies for ballistic-dominated Strombolian-type eruptions at Stromboli, Italy. The latter study is similar to this work in that it used digitized media to quantify eruption dynamics. In terms of analytical procedure, the analysis herein most closely resembles Wilson and Self, who used successive photographs taken at 1-s intervals to estimate upward velocity of a buoyant cloud. This study at Karymsky differs from Wilson and Self in that it measures velocities of the inertial phase of the eruption from incandescent emissions using enhanced time resolution (0.1 s) afforded by modern digital video. [19] Eruption velocity at the base of the camera FOV (V c (h = 0)) is directly measured from the video, starting at time T V, using the ascent rate of incandescent ash and ballistic material rising within the camera s FOV. The spread of incandescence is extracted directly from the 10 frame-per-second (fps) digitized video. At each time step, those pixels surpassing a threshold minimum brightness value are mapped as incandescent. For the Karymsky video imagery, the threshold is set at 25 out of 256 potential gray scale color levels, where background dark pixels from throughout the two nights ranged from 10 to 20. Large, hot ballistics as well as incandescent ash emissions are highlighted in the area of growing incandescence (Figure 9). 8of21

9 Figure 8. Cartoon showing chronology and timing nomenclature that occur during the initiation of typical eruptive events at Karymsky: (a) occurrence of potential precursory seismicity prior to eruptive onset, (b) fragmentation onset that is coincident source of infrasound radiation, (c) appearance of eruption products above the crater rim within the FOVof the camera, (d) arrival of seismic waves at seismic station KRM3, and (e) arrival of acoustic waves at infrasonic microphone at KRM3. This incandescent area tends to grow symmetrically above the initial emission point. After the first few tenths of a second, the ballistics can often be seen to outrace a growing hemiellipse of incandescent pixels corresponding to an expanding cloud of finer particles. This separation is most likely due to lowered susceptibility to drag for denser particles. [20] Measured incremental velocities (during 0.1-s intervals) range up to 65 m/s for the 34 Karymsky 1999 events. In general, these incremental velocity functions are smoothly 9of21

10 in velocity can be observed for the longer time intervals. The velocities calculated at T V and T V s are in fairly good agreement. [21] Velocities calculated from the interval T V! T V s are utilized here as a robust approximation of muzzle velocity, V c (h = 0), because the time interval is short enough to avoid significant biasing from deceleration yet long enough to provide an averaging effect from multiple frames. Calculated values of V c (0) over the course of the first 0.4 s range from 10 to 55 m/s for the suite of 34 eruptive events. This velocity is proportional to the visible plume height achieved during the first 0.4 s, as shown in Figure 11. [22] Assuming that negligible deceleration has occurred within the conduit/vent region, the transit time from initial fragmentation until the plume is first visible is a function of DH, muzzle velocity, and the measured delay time DT A V between acoustic arrival and video of the plume s first appearance. Initial plume appearance time is accurate to within the digitized video frame rate (0 to 0.1 s relative to video frame), which leads to an overall ±0.1-s timing accuracy relative to the GPS time base. The observed values DT A V for the 34 events range from 4.9 to 5.5 s with ±0.1 s accuracy (Table 1). [23] The vertical transit distance within the conduit up until the camera s lower FOV may then be calculated from the following equation: ð DH ¼ D V ADT A V ÞV c V A V c ð2þ Figure 9. (a f) Example sequence of still frames (at 0.1-s intervals) showing the growth of an incandescent eruption cloud above the crater/vent. Right-hand column images are processed as binary colormaps so that areal extent can be calculated at each time step (shown in text inset in left-hand panels). Velocity is calculated from the maximum change in height of the incandescent plume at each time step. decreasing, decelerating slightly during the second-long interval after T V. Table 1 and Figure 10 show velocities calculated during three intervals, T V! T V +0.1s,T V! T V s, and T V! T V s. For many events, a decrease Here V A is approximated as constant velocity (335 m/s) from fragmentation depth up through the conduit length DH and straight line slant distance D to the sensor. Although V A has been measured several times faster within a hot-gasfilled conduit [e.g., Weill et al., 1992; Yokoo and Taniguchi, 2004], the time delay DT A V is insensitive to increased conduit acoustic velocities for open conduit lengths presumed to be a few tens of meters or less. The measured time delay is much more dependent upon conduit length DH and variations in muzzle velocity V c, which is observed to be substantially slower than sound velocity. [24] Minimum slant distance D between crater lip and sensor is extracted from the maximum observed value of DT A V (i.e., 5.5 s). Assuming a conjoint and superficial (DH = 0) source of incandescence and infrasound, D would equal V A DT A V. This implies a slant distance D of at least 1840 m for V A = 335 m/s (i.e., an acoustic velocity at 5 C). The slant distance D in equation (2) is thus set to 1840 m as it provides a minimum distance and still follows within the original estimated range of distances between KRM3 and the vent. [25] Observed values of DT A V range from 4.9 to 5.5 s with mean and median values of 5.3 s (Table 1). Corresponding values of DH according to equation (2) range from 0 to 16 m (Table 1 and Figure 1). Mean and median depths for all 34 events are 6 and 5 m, respectively, but there is a notable bimodal distribution which coincides with the two different acquisition periods (Figure 12). [26] Discussion: Table 1 and Figure 12 provide individual calculation of fragmentation/infrasound generation depths for all 34 events. It is important to remember that these depths are calculated relative to the shallowest event (i.e., 10 of 21

11 Figure 10. Muzzle velocities for 34 events ordered in terms of incandescent rise speeds for T V! T V s averaging interval. Two other curves are calculated from digital video corresponding to T V! T V s and T V! T V +1.0s. DT A V = 5.5 s), which has been arbitrarily forced here to occur at the very top of the conduit due to D being set to 1840 m. On the basis of helicopter overflights of the crater in 1998 (Figure 1), a very shallow fragmentation source seems to be reasonable for at least some events. During these observations, pyroclastic eruptions were seen to be issuing from conduit-choking material only a few meters below the crater rim. Relative depths, as indicated by positive values of DH, imply that fragmentation may occur at variable depths beneath this level. [27] Assuming a relatively stationary depth for the fragmentation source during a sequence of eruptive events, it would be reasonable to expect longer conduit transit times (DT V I ) for material erupted with lower velocities. This relation appears to be borne out in Figure 13, which is a comparison of muzzle velocity and conduit transit time for the 22 events from 11 to 12 September. Solid lines correspond to expected video-acoustic delays for fragmentation depths fixed at 5, 7.5, and 10 m below the camera FOV. Deviation from these trends for the suite of 22 events could be due to a nonstationary fragmentation depth (or potential horizontal position variability) and/or the 0.1-s uncertainty in video timing. Nevertheless, a positive correlation is suggested for muzzle velocity and time delay between acoustic and video onsets. The best fits to the data suggest a source depth 5 10 m below the camera FOV for the night of September. [28] The assumption of conserved rise velocity V c is a simplification that may impact estimates of fragmentation depths. Deceleration because of gravity and drag will overcome inertia of gas, ash, and ballistics to varying degrees within and above the conduit. However, only very grave deceleration will likely result in significantly erroneous shallow fragmentation depths. Deceleration due to gravity will cause V c to diminish by <6 m/s for typical conduit transit times, which are observed to be less than 0.6 s. Gravitational deceleration would then potentially translate to a vertical source depth error of up to 3 m for a few events in the data suite. Deceleration because of drag is slightly more difficult to estimate as the aerodynamic properties (i.e., mass, surface area, and drag coefficients) of the ash and ballistics are not constrained. Smaller less dense par- Figure 11. Evolution of (a) incandescent area occupied by plume and (b) plume height. Calculated muzzle velocity V c (at h = 0) corresponds to rise in incandescent plume height between time T V and T V s. 11 of 21

12 Figure 12. Calculated variations in fragmentation depths (DH) for suite of 34 eruptive events. Shallowest events are arbitrarily set to 0-m depth corresponding to the base of the FOV of the video camera. Event number time axis is arranged chronologically, but is not linear in time. Gap between events #367 and #423 separates the two different nights of video observations. ticles with initial high velocities will decelerate much more quickly because of drag, however, the exact amount of deceleration is difficult to predict. [29] Despite uncertainties related to potential deceleration in the conduit, delay times DT A V between first appearance of incandescence and picked acoustic onsets appear variable and likely reflect minor variations in depth to the fragmentation source. Error bars corresponding to depth uncertainty in Figure 12 are calculated assuming maximum video synchronization and video picking errors of ±0.1 s. Systematic changes in DT A V seem to have occurred between 11 and 12 September and are significant relative to potential timing errors (Figure 12). On average, DT A V decreases to 5.1 s down from 5.4 s, corresponding to an average depth increase of 7 m. Without invoking changes to the source location, this decrease in DT A T would be accommodated by an unreasonable increase in atmospheric sound velocity of 20 m/s (corresponding to a 44 C increase in atmospheric temperature). Time picking errors and camera clock drift are not expected to exceed one frame (0.1 s), so it is most likely that the free-surface level of the fragmenting magma has dropped during the course of 24 hours. [30] Fluctuations in the free-surface level of magma or level of fragmentation within a conduit should not be unanticipated. At Stromboli, the bursting of large gas bubbles at the base of an open conduit is thought to produce infrasound, which has been used to probe for source depth. Assuming a coincident seismoacoustic source at Stromboli, Ripepe and Braun [1994] used seismic and acoustic phases to estimate vertical free-surface fluctuations on the order of 80 m. Another Stromboli study by Ripepe et al. [2002] compared thermal and acoustic signal onsets corroborating magma free-surface variations in excess of tens of meters. Similarly free-surface fluctuations of the basaltic lava lake at Villarrica Volcano have been visually documented and appear to vary by several tens of meters at timescales of weeks to months [Witter et al., 2004]. [31] There have been few studies directed at determining fluid magma levels at more silicic volcanic systems like Karymsky. Visual observations of free-surface levels are often impeded by conduits which appear to be choked with blocks and or tephra so that it is hard to visually ascertain depths. At the andesitic Tungurahua Volcano, Ruiz et al. [2006] attempted to remotely estimate source depth and upward gas velocity and concluded that it was difficult to uniquely determine these parameters with only seismic and acoustic data. Timing of infrasound generation in conjunction with other types of time series data (for example, thermal imagery or video) then provides a promising means to assess evolving source fragmentation depths. Figure 13. The relation between velocities extracted from the video for T V! T V s and video-acoustic time delay DT A V. Timing error bars are ±0.05 s. Solid lines correspond to expected delay times for various depths under the assumption of constant muzzle velocity in the conduit. The 22 displayed events are from the night of September only. 12 of 21

13 4.2. Correlation of Infrasound With Eruption Intensity [32] In addition to there being a reasonable coincidence between the timing of infrasound production and pyroclastic emissions, there is also a clear correlation between infrasound intensity and eruption intensity. As a first-order observation, both the muzzle velocity and plume area expansion, as quantified by digital video, appear to increase with greater peak infrasonic excess pressure [Johnson, 2000]. This relation suggests that infrasound is produced by the explosive upward acceleration of material within the conduit. [33] A quantifiable relation between infrasound and eruptive flux is an important objective at Karymsky and other volcanoes. With proper modeling of recorded infrasound at KRM3, it is possible to recover the initial time history of volume acceleration within the conduit and thereby recover the velocity of erupting material. In the following first-order modeling, several important simplifying assumptions are adopted: [34] 1. It is assumed that the acoustic source behaves as a compact-baffled piston source, with surface dimension significantly smaller than the primary infrasound wavelengths (for example, 225 m for a 1.5-Hz sound). The piston is justified physically as an upward acceleration of material (i.e., gas, ash, and ballistics) within a roughly circular conduit of radius a. For values of a that are small relative to the acoustic wavelength, the source may be considered omnidirectional [Dowling, 1998]. This assumption appears reasonable based upon observations of the crater size during helicopter overflights. [35] 2. Total vertical displacements of this gas/ash/ballistic diaphragm must also be small compared with the wavelength of propagated infrasound, and upward velocity must be subsonic. On the basis of the maximum velocities (<60 m/s) calculated in section 4.1, this is a reasonable approximation; the piston will only rise by several tens of meters during the course of the initial 1.5-Hz infrasound pulse. [36] 3. The dissipative and/or diffractive effects of a fragmentation/infrasound source location that is recessed within the crater/conduit are considered to be small. This assumption may be reasonable considering that calculated values of DH in section 4.1 are less than 20 m, which is significantly less than the primary infrasound wavelengths. [37] 4. Radiation of infrasound is approximated as isotropic and hemispherical. It is assumed that excess pressure falls off as the inverse of distance between source and receiver. Furthermore, intrinsic infrasound attenuation at the distance of KRM3 is negligible and time-invariant. It is further assumed that acoustic site response at the KRM3 sensor is negligible. [38] 5. Pressure sensor instrument response, as measured in a laboratory, has been suitably deconvolved. Processed data are assumed to be flat and calibrated within the bandwidth of interest. [39] For isotropic radiation into a hemisphere, excess pressure (DP) recorded at a slant distance D is then a function of the diaphragm motion [Dowling, 1998] DPðD; t Þ ¼ rpa2 _Ut ð D=V A Þ 2pD ð3þ where _U is the time derivative of the cross-sectional averaged upward velocity, which should be comparable to measured muzzle velocity V c at h = DH. Equation (3) very closely resembles that of a simple acoustic monopole for which the source strength is the time derivative of the volume flux and the far-field term falls off as 1/4pD [Lighthill, 1978]. [40] Solving for the diaphragm-averaged upward velocity in equation (3) gives UT ð I þ tþ ¼ 2D ra 2 ZT Aþt T A DPT ð A þ tþdt ð4þ where the initial starting condition is U(T I ). The integration interval t is set to a short time (ranging from 0.2 to 0.6 s), which corresponds to the entire compressional phase of the recorded infrasound pulse. This is inferred to be the time necessary for the piston diaphragm to reach its maximum upward velocity (U max ). Over this interval, the volumetric flux (U a 2 ) is seen to increase rapidly before peaking (Figure 14). [41] Assuming constant a, the time required to reach maximum velocity is variable for different events and may be a reflection of variable downward-propagation speed of a fragmentation front as postulated by Spieler et al. [2004]. Peak values of volumetric flux (U a 2 ) and timescales t are plotted in Figure 15. There is no apparent relation between integration time interval t and volumetric flux, suggesting that intense high-velocity emissions can be accomplished through either relatively slow or fast piston accelerations. [42] The peak values U max are compared with V c extracted from the digital video as a confirmation of the acoustic waveform modeling. For the 34 analyzed events showcased, U max and V c are in best agreement for a diaphragm radius a of about 3 m. This best-fit radius ranges from 2.7 m for V c calculated during T V! T V s to 3.1 mforv c calculated during T V! T V s (Figure 16). Although the correlation between U max and V c is positive for all video time intervals, the R 2 value for T V! T V s is highest at [43] Discussion: It is important to note that the data correlations presented in Figure 16 are accomplished with a fixed value of a. This value corresponds to an inferred conduit-filling piston radius, which is responsible for uniform upward acceleration of atmosphere. Such a mechanism should be taken as a simplification to the complex flow that exists in the throat of a volcano. As such, a fixed piston radius should not be expected for a sequence of eruptive events in which conduit dimensions and conditions may be changeable. It is likely that conduit diameter, as well as the amount of material choking the conduit, and overall vent geometry will evolve over time, perhaps even from event to event. Such changes may contribute to effective variations in the piston radius. [44] Piston radius a may be solved for individual events so that U max and V c match precisely (Table 1). Utilizing values of V c from the first 0.4 s, a is forced between 1.9 and 4.3 m for all 34 events (with average value of 2.8 m and standard deviation 0.5 m). As estimated from helicopter 13 of 21

14 Figure 14. Time evolution of the scaled volumetric flux (product of upward average velocity and squared conduit radius) for 34 events. Solid line is indicated for the portion of the curve until U max is achieved. overflights, these dimensions appear to be reasonable and changeable depending upon the amount of material choking the crater floor. [45] In general, the fit between material velocity extracted from infrasound and muzzle velocity as measured by video is remarkably good based upon the initial modeling presented above. Most importantly, the positive fit provides strong evidence that the upward acceleration of gases, ash, and ballistics is able to perturb the atmosphere and radiate infrasound to varying degree. This finding presents an alternative to some other models [e.g., Garces and McNutt, 1997; Vergniolle et al., 2004], which suggest that free-surface movements of an intact fluid magma are largely responsible for large-amplitude volcanic infrasound. Although contrasting mechanisms for infrasound generation may still explain recorded infrasound at different volcanic centers (with varying chemistry and eruptive behavior), the explosive expansion of gases is shown here to be a reasonable mechanism that should be commonly considered Coincidence of the Seismoacoustic Source [46] It is presumed here, and in other studies, that explosive fragmentation and/or gas pistoning induce a thrust force upon the volcanic edifice which contributes to a seismic response [e.g., Chouet et al., 1997; Brodsky et al., 1999]. Some of the primary evidence for an explosion response mechanism comes from the temporal association of eruptions with the onset of large-amplitude seismic transients [Lees et al., 2004]. However, the precise relation between timing of the seismic transient and eruptive event is often not well constrained. It is now shown that at Karymsky most events possess some limited low-amplitude precursory seismicity prior to fragmentation (i.e., DT I P > 0) (Figure 8a). Here DT I P is defined as the time difference between inferred infrasonic and seismic source origin times. [47] On the basis of 34 events recorded at KRM3 in 1999, the onset of high-pass-filtered seismicity is observed to occur 3.78 to 8.02 s before the onset of the very impulsive, easily identified infrasound onset. Average values of DT A S are 5.32 s, with a standard deviation 1.14 s (Table 1 and Figure 17). This time delay deviation contrasts with an earlier study at Karymsky where a relatively stationary seismoacoustic delay time (4.06 ± 0.27 s) was determined for a suite of events recorded 1650 m from the vent [Lees et al., 2004]. The different findings may be attributed to a less rigorous search for precursory seismicity performed for the 1997 data set or, alternatively, they may reflect an evolution of the seismoacoustic source mechanism over the intervening two-year period. Certainly the precise determination of seismic initiation is challenging for the 1999 data set. Seismic onset picks are dependent upon the degree of trace data filtering and are subjective because the signals are so emergent. In general, for waveforms recorded at KRM3, there are no clearly identifiable P wave onset times and first motion polarity is wholly impossible to determine. [48] In this study, seismic onset times have been identified only approximately based upon the timing when Figure 15. Volumetric flux (o) and integration time constant (x) for the suite of 34 events. Events are ordered in terms of increasing eruption flux as determined from acoustic traces. 14 of 21

15 Figure 16. Scatterplots comparing U max and V c for fixed vent radii (a) 2.7, (b) 2.8, and (c) 3.1 m and for velocities extracted from video images during time intervals of (a) 0.1, (b) 0.4, and (c) 1.0 s. absolute seismic velocity has exceeded the maximum background seismic level by a factor of two. Such picking is sensitive both to this arbitrary threshold trigger level and to the level of noise in the arbitrary window that is associated with the background. In the case of this analysis, the background noise window is selected as a 3-s interval ending 10 s before the acoustic event begins. Utilizing such values, a composite stack of all 34 events shows DT A S of 5.58 s (Figure 18). The seismic onset is most easily identified by viewing the absolute amplitude of the three-component seismograms in logarithmic amplitude scale. In this manner, amplitude signals are visible relative to background and may be more easily picked. [49] Although the seismoacoustic delay time for the composite stack is not especially sensitive to digital filtering (DT A S remains within ± 0.02 s of 5.58 s), delay times determined for individual events are quite sensitive to the presence of low-frequency energy because of contamination by 0.3-Hz microseisms (Figure 5). In the stacked composite, the microseism noise destructively cancels, but for individual events the microseism noise influences the background seismic level (and subsequent seismic picks) unless it is removed with high-pass filter. Unfiltered data give a range of DT A S from 3.48 to 7.14 s, with a mean delay time of 4.65 s, which is 0.4 s less than that found through the analysis of the high-pass-filtered data. [50] Delay times of 3.80 to 7.86 s for 1-Hz filtered data at KRM3 indicate that seismic radiation is produced within 0 to 4.0 s (±0.2 s) prior to infrasound production. The time duration DT I P of precursory seismicity (i.e., prior to fragmentation) is calculated from the observed time difference DT A S between seismic and infrasonic phases and the expected transit times of each phase: DT I P ¼ DT A S þ D S V S D A V A The estimated ±0.2-s uncertainty in DT I P is calculated from the estimated propagation distances and compressional ð5þ Figure 17. Summary of measured delay times between acoustic and seismic onsets (DT A S, solid circles) and delay times between acoustic and video onsets (DT A V, open circles) for trace data recorded at KRM3. Events are ordered in terms of increasing DT A S. Displayed muzzle velocities (x s), calculated during T V! T V s, are also plotted. 15 of 21

16 Figure 18. (a) Stack of 34 acoustic traces and (c) corresponding vertical-component seismic traces. (b) Absolute amplitudes of the stacked traces and, in the case of the seismic data, normalized logarithmic stack are plotted to emphasize signal onset. wave velocities in the atmosphere (V A ) and ground (V S ). Acoustic velocity (V A ) is fixed at m/s, allowing for atmospheric temperatures to range by a reasonable amount (5 20 C). The corresponding acoustic propagation distance is set to slant distance D plus conduit height DH (1840 < D A < 1860; as determined in section 4.1). Seismic P wave propagation distances are not precisely known because of poor seismic depth constraints. Here they are conservatively assumed to be somewhat less than the slant distance D and greater than the horizontal distance between vent and station (1700 < D S < 1840). Maximum seismic P wave velocity is taken as 1200 m/s based upon the apparent velocities recovered from a radial array of seismic sensors deployed 1.6 km from the vent in 1998 [Johnson et al., 2003]. Using these parameters, DT A S time delays for a coincident seismoacoustic source would be expected to range between 4.14 and 3.82 s (or 3.98 ± 0.16 s). [51] Discussion: Although there is no evidence for intense preeruption short-period seismicity at Karymsky in 1999, emergent low-intensity seismicity precedes fragmentation (by up to 4 s) and appears common for the majority events. This energy comprises a very small fraction of the total seismic yield but is significant because it directly leads up to the explosive event onset. For these reasons, and largely because the seismic precursor envelope appears to ramp up to the eruption onset, it is more likely that the early seismicity reflects a near-vent conduit priming phenomenon rather than a spatially distinct volcanic earthquake that might trigger a seismic event in the fluid upper conduit [e.g., Sturtevant et al., 1996]. [52] Further source constraints related to the location and mechanism of the precursory seismicity are difficult to extract given the existing seismic data, but it is possible to speculate on potential source motions. Among several possibilities, it is conceivable that the precursory seismicity reflects a slow initiation of magma foam collapse. It is also possible that preexplosion magma deformation due to bubble expansion, bubble film failure, or generalized bubble coalescence may induce stress that is able to radiate seismically. Alternatively, it could be a response to fracturing of caprock as pressurized gases make their way to the surface. In this scenario, the seismic onset might be coincident with initial magma fragmentation and the infrasound onset would be coincident with explosive gas expansion occurring at the free surface. It is interesting to note that there is no clear correlation between the amount of precursory seismicity and the apparent depth of the fragmentation source and/or the initial intensity of the eruption (Figure 17). [53] At several other volcanoes, there is clear evidence to support that seismicity occurs prior to, and leads up to, infrasound generation. At both the dacitic Guagua Pichincha in 1999 [Johnson et al., 2003] and dacitic Mount St. Helens in 2005 [Moran et al., 2004], seismicity on occasion appears to precede the primary eruptive phases by tens of seconds. It has been speculated that such precursory seismicity is due to fracturing of conduit rocks during the flow of pressurized gas en route to the surface vent [Johnson 16 of 21