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Journal of Volcanology and Geothermal Research 137 (2004) 219–230

Seismoacoustic measurements during the July–August 2001

eruption of Mt. Etna volcano, Italy

S. Grestaa,*, M. Ripepeb, E. Marchettib, S. D’Amicoc, M. Coltellic,A.J.L. Harrisd, E. Priviterac

aDipartimento di Scienze Geologiche, Universita di Catania, Corso Italia 55, 95129 Catania, ItalybDipartimento di Scienze della Terra, Universita di Firenze, Italy

c Istituto Nazionale di Geofisica e Vulcanologia, Sezione di Catania, ItalydHIGP/SOEST, University of Hawaii, Honolulu, HI, USA

Abstract

On July 18, 2001, two main eruptive vents opened on the southern flank of Mount Etna volcano (Italy) at f2100 m and

f 2550 m a.s.l., respectively. The former vent fed mild strombolian activity and lava flows, while the latter represented the

main explosive vent, producing strong phreato-magmatic explosions. Explosions at this latter vent, however, shifted to a

strombolian style in the following days, before switching back to phreato-magmatic activity towards the end of the eruption,

which ended on August 9, 2001.

On August 3, a small seismoacoustic array was deployed close to the eruptive vents. The array was composed of three

stations, which recorded seismic and infrasonic waves coming from both of the eruptive vents. A further seismoacoustic station,

equipped with a thermal-infrared sensor, was also installed several kilometers north of the first array. Seismic signals relating to

the strombolian activity at the 2100-m vent were characterized by a strong decompression at the source. Analysis of the time

delays between seismic, infrasonic and infrared event onsets also revealed that ejection velocities during explosions from both

vents were subsonic. Time delays between the onset of explosive events apparent in the infrared and infrasound data indicated

that the explosion source at the 2550-m vent was located 220–250 m below the crater rim. In comparison, the depth of the

seismic source was estimated to be between 230 and 335 m below the rim. This converts to 120–150 and 130–235 m below the

preexisting ground surface. In addition, time delays between seismic and infrasonic signals recorded for the lower (2100 m)

vent also revealed a seismic source that was no more than a few tens of meters deeper than the fragmentation surface.

D 2004 Elsevier B.V. All rights reserved.

Keywords: Mt. Etna; explosive eruptions; arrays; seismic; infrasonic; thermal data

1. Introduction

The volcanic activity of Mount Etna volcano

(Italy) may be roughly divided into two main types:

0377-0273/$ - see front matter D 2004 Elsevier B.V. All rights reserved.

doi:10.1016/j.jvolgeores.2004.05.017

* Corresponding author. Tel.: +39-095-7195709; fax: +39-095-

7195728.

E-mail address: gresta@unict.it (S. Gresta).

flank eruptions and persistent activity at the summit

craters. While the former is mainly characterized by

the opening of vents and/or fractures to feed lava flow

on the flanks of the volcano, the latter is characterized

by many different, and sometimes coexisting, styles of

summit-located activity. These styles of activity in-

clude persistent degassing, strombolian or phreato-

magmatic explosions, lava fountaining, crater filling

S. Gresta et al. / Journal of Volcanology and Geothermal Research 137 (2004) 219–230220

or collapse events, and low effusion rate lava emis-

sions (Cristofolini et al., 1988; Harris et al., 2000).

Volcanic tremor and long-period (LP) events are

generally associated with eruptive processes induced

by summit degassing (e.g. Aki et al., 1977; Chouet,

1988) and are observed on many persistently active

volcanoes throughout the world. Such events are

evident at Etna, where tremor has been related to

gas escaping through the central conduits (Schick and

Riuscetti, 1973). In this regard, peaks in spectral

frequency recorded at Etna can be related to the

magma–gas mixture rising in the central conduit

(Riuscetti et al., 1977; Seidl et al., 1981). In addition,

two different tremor sources have been proposed at

Etna. The former is associated with processes occur-

ring in the upper portion of the central conduits,

whereas the latter originates in a flat magma reservoir,

located ca. 2 km beneath the summit of the volcano

(Schick et al., 1982). Further investigation has shown

that during persistent degassing, the tremor source is

confined to the shallowest, i.e. upper part, of the

central conduits (Del Pezzo et al., 1993). However,

during paroxysmal summit eruptions, deeper sources

also show seismic activity (e.g. Ferrucci et al., 1990;

Gresta et al., 1996; Privitera et al., 2003). Finally,

studies of tremor amplitude at different stations across

Etna have shown that, when a flank eruption occurs,

the main tremor source is almost coincident with the

ephemeral flank vent (e.g. Falsaperla et al., 1994).

Indeed, the tremor signal always appears as a contin-

uous vibration at seismic stations located some kilo-

meters from the active vents. At permanent stations

closest to the active summit crater(s), tremor some-

times appears as a continuous vibration with an almost

constant amplitude, but sometimes with many small

impulsive events (Gresta et al., 1987).

Studies that have examined the LP events (e.g.

Chouet et al., 1994; Falsaperla et al., 2002) have

shown a strong similarity between their spectral con-

tents and those obtained for volcanic tremor. It was

thus concluded that at Etna, as at other volcanoes (e.g.

Fehler and Chouet, 1982; Malone, 1983; Chouet,

1988), tremor and LP events originate from similar

physical process occurring in approximately the same

location. Attempts to model and locate the tremor

(and/or LP events) source at Mt. Etna must take into

account the high heterogeneity of the region where

active structural features, shallow and deep magma

bodies, dyke intrusions, etc., are present. This factor

may produce significant variations in the features of

the seismic signals collected at the various stations of

permanent networks (e.g. Gordeev, 1993).

New insights into explosive volcanic processes

have been achieved in recent years using infrasonic

data (Vergniolle and Brandeis, 1994; Buckingham and

Garces, 1996; Ripepe et al., 1996). Analysis of

seismic records combined with infrasonic signals

has revealed that shallow volcanic tremor can be

related to magma degassing (Ripepe et al., 1996;

Johnson et al., 1998). This process appears on infra-

sonic records as low-pressure pulses generated at a 1–

2 s rate and have been interpreted as the bursting of

small gas pockets at the magma free surface (Verg-

niolle and Brandeis, 1996; Ripepe and Gordeev,

1999).

During a previous seismoacoustic experiment per-

formed at Etna, a polarization analysis of the seismic

wave field recorded at a station located 50 m from the

active crater rim revealed a large content of P-waves

(Ripepe et al., 2001a). The volcanic tremor had an

amplitude modulation that correlated (at a rate of

about 2 s) with acoustic pressure bursts (from 1 to 5

Pa). In addition, the infrasonic wave field was coher-

ent over several hundreds of meters, whereas the

tremor wave field rapidly lost its character at shorter

distances. Seismic and infrasonic wave fields were

quite well correlated, indicating that the same source

process was responsible for both the volcanic tremor

and the infrasonic bursts. Moreover, the short time

lags (ca. 0.4 s) between volcanic tremor and infrasonic

bursts supported a shallow origin of both volcanic

tremor and infrasonic pulses. The conclusion was that

volcanic tremor recorded at Etna during a period of

vigorous strombolian activity at the summit craters

was generated by the superposition of small, shallow,

point sources: a consequence of degassing at the top

of the magma column (Ripepe et al., 2001a).

In this paper, we analyze seismic and acoustic

signals collected at Etna during the last stages of the

July–August 2001 flank eruption. Data were col-

lected using a small array deployed close to two

highly explosive flank vents. The study also benefit-

ed from data provided by a thermal-infrared sensor.

This was the first time that such a multidisciplinary

experiment was performed during a flank eruption at

Etna, where the aim of our work was to characterize

S. Gresta et al. / Journal of Volcanology and Geothermal Research 137 (2004) 219–230 221

the explosion source and to calculate both the depth

of the explosions in the feeder conduit and the

velocity of the ejection. In this regard, the prime

function of this paper is to extend the approach of

Ripepe et al. (2002) and show how a carefully

designed monitoring system based on seismoacous-

tic and thermal data can yield reliable quantitative

parameters for explosive volcanic eruptions. In car-

rying out such an approach, waveform analysis and

signal amplitudes are not important, but instead,

signal onset times become the focus of the analysis.

These times allow us to extract locational data for

the source of explosions. In doing this, we generate

quantitative parameters that constrain the location

and depth of explosion sources during the 2001

eruption of Etna.

Fig. 1. Sketch map of the southern flank of Mt. Etna. The location

of both the main eruptive vents and recording stations (white dots)

considered in this study are marked.

2. Volcanic activity

Prior to the 2001 flank eruption, the previous

flank eruption of Mt. Etna took place between

December 1991 and March 1993 (e.g. Calvari et

al., 1994). Volumetrically, the 1991–1993 eruption

was the largest effusive eruption to occur at Etna

during the last three centuries. After this event, the

eruptive activity was confined to the summit craters

(the Northeast Crater, Bocca Nuova, the Voragine

and Southeast Crater) until 2001. Between 1995 and

1998, activity occurred mainly at Northeast Crater

and Bocca Nuova. However, a sharp increase in

activity began during 1998 with a series of 23 lava

fountains taking place at the Voragine and Southeast

Craters between June 1998 and February 1999.

During this 9-month-long eruptive period, two lava

flows were also erupted from the Bocca Nuova and

Southeast Crater (e.g. Calvari et al., 2002; Harris

and Neri, 2002). Activity then became concentrated

at Southeast Crater with a series 66 lava fountain

episodes occurring between January and June 2000.

In addition, lava flow emission from Southeast

Crater began during January 2001 and another series

of lava fountains began on May 9, 2001. The

duration of each lava fountain episode generally

ranged from few tens of minutes to a few hours.

During the early morning of July 13, 2001, the

17th paroxysmal eruption of the final fountain series

occurred at Southeast Crater (Fig. 1). It was charac-

terized by strong strombolian activity with lava foun-

taining and lava overflows. About 800 earthquakes

occurred in the 15 h encompassing this short-lived

eruption. During the following 6 days, more than

2500 earthquakes (1VMdV 3.9) were recorded, main-

ly on the southern flank of the volcano. At the same

time, a complex fracture field opened on the upper

slopes of the volcano, extending from the base of

Southeast Crater. This fracture system became erup-

tive on July 17 (for details, see Research Staff of

INGV, 2001).

During the early morning of July 18, a further

flurry of seismic events accompanied the opening of a

second eruptive vent at about 2100 m a.s.l. (see Fig.

Fig. 2. Picture of the 2100 Crater showing explosive and lava flow

activity from the cone base.

Fig. 3. Picture of the 2550 Crater showing its typical activity. The whole c

several kilometers and ash fall hindered the use of Catania airport for som

S. Gresta et al. / Journal of Volcanology and Geothermal Research 137 (2004) 219–230222

1). This vent (described hereafter as ‘‘2100 Crater’’)

fed mild strombolian activity from two sources at the

top of its cone and lava flows from a fracture at the

cone base (see Fig. 2). On the same day, a further

eruptive vent opened at about 2550 m a.s.l. (Fig. 1).

This last vent (hereafter termed ‘‘2550 Crater’’) was

the main explosive vent, producing powerful phreato-

magmatic explosions with high ash columns (Fig. 3).

However, over the following days, activity became

characterized by strombolian activity and lava foun-

tains, before returning again to a phreato-magmatic

explosive style in the closing stages of the eruption.

The eruptive activity (both explosive and effusive)

continued vigorously until the first days of August at

both the 2100 and 2550 Craters, making any close

approach extremely dangerous. During August 3–4,

explosive activity at the 2550 Crater was continuing,

with emission of both ash and incandescent ejecta at

an average rate of three explosions every hour. At this

point, the elevation of the ash column was about 1.5–

2 km, and the output rate of ejecta was estimated,

from Doppler radar measurements, to be about 25000

one was built in a few days; the ash column reached an elevation of

e days.

Fig. 4. An example of east–west ground displacement and pressure

transient records (CCZ station) produced by an explosion at 2100

Crater. The time delay (shaded area) between the onset of the two

signals remained stable at 1.54 s for the duration of the study period.

S. Gresta et al. / Journal of Volcanology and Geothermal Research 137 (2004) 219–230 223

kg/s. At the same time, the 2100 Crater emitted both

bombs (to a distance of a few tens of meters from the

crater rim) and ash. Towards August 5, the rate of lava

output sharply declined, and ash composition showed

some significant variation forerunning the end of the

eruption (Taddeucci et al., 2002).

Fig. 5. An example of seismic, thermal and acoustic signals related to an

onsets are ca. 3.8 s (dark shaded area), and thermal-infrasound time delays

studied period.

On August 5, explosive activity at the 2550 Crater

was continuing with ash explosions (up to three per

minute) feeding a f 1-km-high ash plume. The 2100

Crater, however, showed a decrease in activity from

moderate to sporadic emission of ejecta. On August 6,

the explosive activity at the 2550 Crater stopped and

the 2100 Crater showed only sporadic strombolian

activity, which ended on the following day. Lava

effusion completely stopped late on August 9.

3. Data acquisition

On August 3, a seismoacoustic array, composed

of three stations, was deployed close to the 2100

Crater (see Fig. 1). These stations, CCZ, MSI and

MEM, were located 320, 1220 and 1610 m from

the 2100 Crater, respectively. Each station was

equipped with a three-component Lennartz seis-

mometer (T= 5 s; sensitivity 160 V/m/s) and acous-

tic sensors. Each infrasonic sensor consisted of a

Monacor condenser microphone MC-2005, with a

sensitivity of 46 mV/Pa in the 1–20 Hz infrasonic

band. Data were recorded using a five-channel 16-

bit digital acquisition system at a sampling rate of

explosion at 2550 Crater. Time delays between seismic and thermal

are 2.7 s (light shaded area). Both were stable for the duration of the

S. Gresta et al. / Journal of Volcanology and Geothermal Research 137 (2004) 219–230224

18.5 ms, with a 20-Hz anti-aliasing filter. Time

synchronization was achieved by recording the

DCF radio code.

A fourth station was installed at the Torre del

Filosofo (TDF) site (Fig. 1). Seismometers and acous-

tic sensors at this station were of the same type as at

the first stations, but the TDF station was also

equipped with an Omega thermal-infrared thermome-

ter. This sensor operates in the 8- and 14-Am range

and had the same acquisition rate (18.5 ms) as the

whole system.

Each station ran continuously until the end of the

eruption, where Fig. 4 shows an example of the seismic

and infrasonic data collected. In addition, an example

of seismic, thermal and acoustic data related to an

explosion at the 2550 Crater is shown in Fig. 5.

Fig. 6. Example of the particle motion typically recorded at the three

stations of the array during an explosion at the 2100 Crater.

Polarization towards the explosive vent is evident only at the nearest

station (CCZ).

4. Analysis of data

4.1. Features of the seismic signals

Fig. 6 shows the particle motion typically

recorded at the three stations of the 2100-m array

during an explosion at the 2100 Crater. The polar-

ization towards the vent is evident only at the

nearest station (CCZ), located 320 m from the

active vent. At the two stations MSI and MEM,

located ca. 2 km from the eruptive vent, the

polarization is lost. This is consistent with previous

results obtained for the summit craters of Mt. Etna,

where the tremor wave field rapidly loses its fea-

tures at distances of a few hundreds of meters from

the source (Ripepe et al., 2001a).

Fig. 4 shows an example of the seismic displace-

ment, produced by the explosive activity of the

2100 Crater, recorded at the CCZ station. The EW

component of the displacement indicates an initial

contraction of the source. This seismic waveform

was stable in time reflecting a stable source mech-

anism. This is consistent with visual observations

that showed that eruptive activity remained constant

for the whole period when the array was operating.

In addition, contraction of the source has been

commonly observed at Stromboli (Neuberg et al.,

1994; Chouet et al., 1999), Aso (Kaneshima et al.,

1996), Erebus (Rowe et al., 2000) and Sakurajima

(Uhira and Takeo, 1994). This, however, is the first

time that decompression of the source has been

clearly observed at Mt. Etna. Such a contracting

source during explosive activity is predicted by the

‘‘foam collapse model’’ (Jaupart and Vergniolle,

1988), where Ripepe and Gordeev (1999) have

demonstrated that an explosion may be triggered

by a pressure drop in the conduit, which is seismi-

cally recorded as a contraction in the area close to

the source. As soon as the foam collapses, a large

gas pocket (or bubble) expands in the conduit and

magma is forced to flow downward around the

bubble. Laboratory experiments show that a decom-

pressive seismic signal is generated during the gas

expansion as a consequence of the foam collapse

(Ripepe et al., 2001b).

S. Gresta et al. / Journal of Volcanology and Geothermal Research 137 (2004) 219–230 225

4.2. Location of the seismoacoustic source

The infrasonic onsets at the three stations of the

array have been used to calculate the direction of the

sonic beam. This technique allows us to determine the

direction from which the sound is propagating and is a

simple but effective method to monitor the evolution of

the activity at different craters. In our case, a beam-

forming technique cannot be applied to locate the

source of the infrasonic waves recorded by the array.

Beam forming is based on the assumption that the wave

front is planar and aims to identify the azimuth and

velocity of a plane wave that propagates across the

array. In our case, the distance of the stations from the

source (350–3000 m) is comparable (or larger) to the

size of the array (ca. 1300 m). In this situation, the

acoustic wave-front cannot be considered planar. We

have thus applied a semblance method based on a grid

search procedure, where we search for the location of

the highest probability that a single point in three-

dimensional space has to be the source. Semblance is

calculated between the theoretical and measured dis-

tribution and is assumed to be representative of the

probability that a single node is part of the irradiating

source volume (for details, see Ripepe and Marchetti,

2002).

We located the acoustic source using this grid

searching analysis by considering a 100-m equi-spaced

Fig. 7. Percentage distribution (from 50% to 100%) of the semblance functi

the 2550 Crater (right). The source is located in the node that has the hi

highest beam amplitude (for details, see Ripepe and Marchetti, 2002).

grid over an area of 5� 7 km and then computed the

theoretical time delays for each node. Theoretical delay

times were then used to delay-and-sum the infrasonic

waveforms. All computations were performed using a

reference plane, the elevation of which was 2100 m

a.s.l. Acoustic signals produced by the explosions

show two recurrent back-azimuths, at N315j and

N340j. These are consistent with the direction of the

2100 and 2550 Craters, respectively (Fig. 7). This

result shows the reliability of the location method as

well as the high degree of resolution with which we can

locate the source of the acoustic waves. Moreover,

acoustic signals propagate from the source to the

sensors with almost no path distortion and can thus

be directly assumed to be representative of the source

waveform. This simplicity in the acoustic wave has

been used to discriminate different sources and to

reveal changes in the dynamics of the source at Strom-

boli volcano (Ripepe and Marchetti, 2002).

Identification of the source location allowed us to

separate infrasonic signals produced by the two

active craters. For each explosion, we first delayed

the signals recorded at the array and then, after a

correction for geometrical spreading, stacked them.

In this way, we reduced possible topographic reflec-

tions around the station and noise produced by

secondary sources, such as wind. Acoustic signals

produced by both the 2100 and the 2550 Crater (see

on used to calculate the sonic beam for the 2100 Crater (left) and for

ghest semblance with the measured delay time distribution and the

Fig. 8. Acoustic waveforms recorded at the three stations of the array relating to explosions at the 2100 Crater (upper signals) and at the 2550

Crater (lower signals).

S. Gresta et al. / Journal of Volcanology and Geothermal Research 137 (2004) 219–230226

Fig. 8) are not long-lasting signals, as modelled by

Buckingham and Garces (1996). This excludes con-

duit resonance as their source mechanism.

4.3. Source process and depth of the explosions in the

magma conduits

For the 2550 Crater, the time delay between the

thermal and the infrasonic signals (Dte) recorded at

TDF station was rather stable (ca. 2.7 s) during the

whole experiment. In comparison, the time delay

between the thermal and the seismic onsets (Dtr) was

around 3.8 s.

The time delay between the thermal and infra-

sonic onsets can be used to estimate the fragmen-

tation depth in the conduit (Ripepe et al., 2002). As

sketched in Fig. 9, the thermal signal for the same

event will travel up the conduit length (h) at the

speed of the ascending gas–pyroclast mixture (Ujet).

The signal will then cover the distance (x) to the

sensor at the speed of light. Thus, at the sensor-to-

vent distances considered here, the time required for

the signal to travel distance x is essentially zero.

Hence, the time that the thermal signal takes to

travel from the source to the receiver (te) reduces to

te = h/Ujet. The delay between the thermal and the

infrasonic signal (Dte), or how long the thermal

signal (te) lags behind the infrasonic signal (ti),

can thus be defined as:

Dte ¼ ðti � teÞ ¼ hðUjet � cVÞUjetcV

þ x

c

� �ð1Þ

where c and cV are the speed of sound in air and in

the conduit, respectively; h is the depth of the

Fig. 9. Simple sketch showing the estimation of the source depth(s)

using seismic, acoustic and/or thermal data at a given station. See

text for symbol definition.

S. Gresta et al. / Journal of Volcanology and Geothermal Research 137 (2004) 219–230 227

magma free surface (or fragmentation surface) and

Ujet is the gas jet velocity (see Fig. 9).

Eq. (1) can thus be used to infer changes in the

depth of the magma free surface (h) and/or the gas jet

velocity (Ripepe et al., 2002). Given a hypothetical

explosion from a source that is level with the rim of

the 2550 Crater (i.e. h = 0), the infrasonic signal

should arrive at TDF station (xf 1760 m) approxi-

mately 5.17 s after the thermal signal. Therefore, the

observed time delays (Dte) of around 2.7 s (signifi-

cantly less than 5.17 s) indicate that the infrasound is

generated within the conduit, and that the gas jet

velocity is subsonic (Ujet < c V). We assume that

c = 340 m/s and Ujet = 80 m/s (Dubosclard et al.,

1999) and that cV ranges between 708 m/s, if travel

time in a pure hot gas is considered (Weill et al.,

1992), and 340 m/s, the sound speed in air. Given

these assumptions, the estimated depth of the magma

free surface within the conduit of the 2550 Crater

ranges from 220 to 250 m below the crater rim.

Conversely, the time delay (Dtr) between the

seismic and infrasonic onsets may be used to define

the depth (H) of the explosion in the conduit (Fig. 9)

with respect to the magma free surface (Ripepe et al.,

2001b) from:

Dtr ¼ ðte � tsÞ ¼H

Ub

� �þ h

Ujet

� �

� ½x2 þ ðhþ HÞ2�1=2

að2Þ

Here, all parameters are the same as in Eq. (1),

except for a, which is the seismic (P-wave) velocity,

and Ub, which is the velocity of gas moving through

the magma. We used a = 1800 m/s (see Privitera et al.,

2003), with two values (5 and 50 m/s) for Ub. These

values are taken to be reasonable values for the

buoyancy condition of the bubble and for the accel-

erated expansion of a gas bubble rising in a vertical

tube (Wallis, 1969). The computed H values are 8–10

and 80–100 m, respectively. Therefore, the depth

(h +H) of the seismic source from the 2550 Crater

rim ranges between 230 and 335 m.

Measurements were performed during the final

phases of the eruption, when the eruptive cone had

its final elevation of about 100 m; this implies that a

depth below the previous ground surface was ap-

proximately 100–150 m for the magma free surface

and no more than 250 m for the seismic source. At

this time, the eruptive activity was phreato-magmatic

in character, and these depths compare with a depth

of 400–600 m obtained for the local water-table in

this location (Pham et al., 1982).

Due to the relative position between the stations

and the eruptive craters (see Fig. 1), no thermal data

are available for the 2100 Crater. Fig. 4 shows the

typical record of the seismic and infrasonic signal

due to an explosion at 2100 Crater and recorded at

the CCZ station. The time delay between the

seismic and infrasonic onsets (Dt) was very stable

at 1.54 s (Fig. 4) over the whole experiment. This

stability of the time delay strongly supports the

hypothesis of a stationary explosion source and

has been previously observed at Karymsky volcano

(Johnson et al., 1998), but not at Stromboli (Ripepe

and Braun, 1994) or Mt. Erebus (Aster et al., 2003).

Also, the explosion depth may be estimated by

considering the distance x between the station

CCZ and the active vent (c 320 m). Now the time

delay (Dt) between the infrasonic and the seismic

S. Gresta et al. / Journal of Volcanology and Geothermal Research 137 (2004) 219–230228

onsets may be used to define the depth of the

explosion in the conduit (Ripepe et al., 2001b) as

follows

Dt ¼ ðti � tsÞ ¼H

Ub

� �þ h

cV

� �þ x

c

� �

� ½x2 þ ðhþ HÞ2�1=2

að3Þ

Here, symbols are the same as in Eq. (2). We first

assume that seismic and infrasound pulses are both

simultaneously generated close to the free surface of

the magma column. This is the boundary condition

(H = 0) allowing us to estimate the largest possible

depth of the explosion source. The computed depth

(h) ranges between 300 and 740 m. As a second

step, we consider that the seismic and infrasound

sources are not coincident. We may then estimate

the possible values of the depth (H) of the explosive

source from the magma free surface (see Fig. 9).

For our first buoyancy condition (Ub = 5 m/s), H

cannot exceed 4 m. In the case of accelerated gas

expansion (Ub = 50 m/s), H is less than 40 m. Thus,

the depth of the seismic source relative to the

magma free surface in the conduit of the 2100

Crater may range from a few meters to a few tens

of meters. However, because of our lack of thermal

data, no absolute values for the depth of the magma

free surface from the crater rim can be obtained.

5. Concluding remarks

A short multiparameter geophysical experiment

was performed at Mt. Etna volcano during the closing

phases of the July–August 2001 flank eruption. A

three-station seismoacoustic array was deployed close

to the eruptive vents along with a fourth station that

also provided thermal data. Analysis of these integrat-

ed data sets yielded some interesting results that can

be summarized as follows:

(1) The seismic wave-field loses its character at

distances of more than a few hundreds of meters

from the source, confirming previous results

obtained at Mt. Etna (Ripepe et al., 2001a). In

fact, the polarization towards the eruptive vent

was clearly evident only in data obtained at the

nearest station to the active crater.

(2) The seismic waveform was stable in time and

indicated decompression of the source. The

‘‘foam collapse model’’ was invoked as the

source, with the explosion triggered by a pressure

drop inside the conduit.

(3) Our seismoacoustic investigation allowed us to

locate the explosive source and to discriminate,

with reliability, between the two eruptive craters.

This kind of investigation may be a powerful tool

in identification of the location of explosive vents

in poor weather conditions and to discriminate

the active vent in cases where multiple vents are

present.

(4) No evidence of temporal variation in mecha-

nism, location or depth of the sources was

found.

(5) The combined use of seismic, acoustic and

thermal sensors allowed us to define the depth

(below the preexisting ground surface) of both

the explosions (100–150 m) and the seismic

source (f 250 m). These results are in

agreement with a shallow, phreato-magmatic

origin of the explosive activity at the 2550

crater.

(6) The lack of thermal data for the 2100 Crater

hindered the estimate of the depth of the magma

free surface at this location. However, the depth

of the seismic source relative to the magma free

surface was estimated to be in the range of a few

meters to a few tens of meters.

Results reported in points (5) and (6) represent

the only quantitative constrains of the source mech-

anism of the last stages of the July–August 2001

eruption at Mt. Etna volcano. Moreover, our results

show that a carefully designed monitoring system

based on seismoacoustic and thermal data can yield

reliable quantitative parameters for explosive volca-

nic eruptions. The experiment discussed here can

thus be considered field test for a future monitoring

system at Mount Etna volcano.

Acknowledgements

Detailed information regarding eruptive activity has

been taken from the daily reports published on the Web

by the Istituto Nazionale di Geofisica e Vulcanologia.

S. Gresta et al. / Journal of Volcanology and Geothermal Research 137 (2004) 219–230 229

Pictures have been taken from the Web site http://

www.strombolionline.com.We acknowledgeMassimo

‘‘Fioba’’ Della Schiava for field help during the

deployment of the array and data collection. We thank

Jackie Caplan-Auerbach and Sylvie Vergniolle for their

stimulating and careful reviews of this paper.

This research has been funded by Italian National

Group of Volcanology (INGV-GNV) and by NSF

grant EAR-0207734.

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