Post on 27-Apr-2023
www.elsevier.com/locate/jvolgeores
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.
References
Aki, K., Fehler, M., Das, S., 1977. Source mechanism of volcanic
tremor: fluid driven crack models and their application to the
1963 Kilauea eruption. J. Volcanol. Geotherm. Res. 2, 259–287.
Aster, R., Mah, S.Y., Kyle, P., McIntosh, W., Dunbar, N., Johnson,
J., Ruiz, M., McNamara, S., 2003. Very long period oscillations
of Mount Erebus Volcano. J. Geophys. Res. 108, 2522
(doi:10.1029/2002JB002101).
Buckingham, M.J., Garces, M.A., 1996. Canonical model of vol-
cano acoustics. J. Geophys. Res. 101, 8129–8151.
Calvari, S., Coltelli, M., Neri, M., Pompilio, M., Scribano, V., 1994.
The 1991–1993 Etna eruption: chronology and lava flow-field
evolution. Acta Vulcanol. 4, 1–14.
Calvari, S., Neri, M., Pinkerton, H., 2002. Effusion rate estimations
during the 1999 summit eruptions on Mount Etna, and growth
of two distinct lava flow fields. J. Volcanol. Geotherm. Res. 119,
107–123.
Chouet, B., 1988. Resonance of a fluid-driven crack: radiation
properties and implications for the source of long-period events
and harmonic tremor. J. Geophys. Res. 93, 4375–4400.
Chouet, B., Dawson, P., Falsaperla, S., Privitera, E., 1994. A char-
acterization of long period events recorded during the eruptive
activity of Mt. Etna, Italy, in 1992. Acta Vulcanol. 4, 81–86.
Chouet, B., Saccorotti, G., Dawson, P., Martini, M., Scarpa, R., De
Luca, G., Milana, G., Cattaneo, M., 1999. Broadband measure-
ments of the sources of explosions at Stromboli volcano, Italy.
Geophys. Res. Lett. 28, 1937–1940.
Cristofolini, R., Gresta, S., Imposa, S., Patane, G., 1988. Feed-
ing mechanism of eruptive activity at Mt. Etna based on
seismological and petrological data. In: King, C.Y., Scarpa,
R. (Eds.), Modeling of Volcanic Processes. Springer, Berlin,
pp. 73–93.
Del Pezzo, E., De Martino, S., Gresta, S., Martini, M., Milana, G.,
Patane, D., Sabbarese, C., 1993. Velocity and spectral character-
istics of the volcanic tremor at Etna deduced by a small seis-
mometer array. J. Volcanol. Geotherm. Res. 56, 369–378.
Dubosclard, G., Cordesses, R., Allard, P., Hervier, C., Coltelli,
M., Kornprobst, J., 1999. First testing of a volcano Doppler
radar (Voldorad) at Mount Etna, Italy. Geophys. Res. Lett. 26,
3389–3392.
Falsaperla, S., Privitera, E., Spampinato, S., Cardaci, C., 1994.
Seismic activity and volcanic tremor related to the December
14, 1991 Mt. Etna eruption. Acta Vulcanol. 4, 63–73.
Falsaperla, S., Privitera, E., Chouet, B., Dawson, P., 2002. Analysis
of long period events recorded at Mt. Etna (Italy) in 1992 and
their relationship to eruptive activity. J. Volcanol. Geotherm.
Res. 114, 419–440.
Fehler, M., Chouet, B., 1982. Operation of a digital seismic net-
work on Mount St. Helens volcano and observations of long-
period seismic events that originate under the volcano. Geophys.
Res. Lett. 9, 1017–1020.
Ferrucci, F., Godano, C., Pino, N.A., 1990. Approach to the volca-
nic tremor by covariance analysis: application to the 1989 erup-
tion of Mt. Etna (Sicily). Geophys. Res. Lett. 17, 2425–2428.
Gresta, S., Imposa, S., Patane, D., Patane, G., 1987. Volcanic trem-
or at Mt. Etna: state of the art and perspectives. Pure Appl.
Geophys. 125, 1079–1095.
Gresta, S., Lombardo, G., Cristofolini, R., 1996. Characteristics
of volcanic tremor accompanying the September 24th, 1986
explosive eruption of Mt. Etna (Italy). Ann. Geofis. 39,
411–420.
Gordeev, E., 1993. Modeling of volcanic tremor as explosive point
sources in a singled-layered, elastic half-space. J. Geophys. Res.
98, 19687–19703.
Harris, A.J.L., Neri, M., 2002. Volumetric observations during par-
oxysmal eruptions at Mount Etna: pressurized drainage of a
shallow chamber or pulsed supply? J. Volcanol. Geotherm.
Res. 116, 79–95.
Harris, A.J.L., Murray, J.B., Aries, S.E., Davies, M.A., Flynn, L.P.,
Wooster, M.J., Wright, R., Rothery, D.A., 2000. Effusion rate
trends at Etna and Krafla and their implications for eruptive
mechanisms. J. Volcanol. Geotherm. Res. 102, 237–269.
Jaupart, C., Vergniolle, S., 1988. Laboratory models of Hawaiian
and Strombolian eruptions. Nature 331, 58–60.
Johnson, J., Lees, J., Gordeev, E., 1998. Degassing explosions at
Karymsky volcano, Kamchatka. Geophys. Res. Lett. 25,
3995–3998.
Kaneshima, S., Kawakatsu, H., Matsubayashi, H., Sudo, Y., Tsut-
sui, T., Ohminato, T., Ito, H., Uhira, K., Yamasato, H., Oikawa,
J., Takeo, M., Iidaka, T., 1996. Eruption process of Aso volcano
inferred from near-field broad-band seismic observations. Sci-
ence 273, 642–645.
Malone, S.D., 1983. Volcanic earthquakes: examples from Mt. St.
Helens. In: Kanamori, H., Boschi, E. (Eds.), Earthquakes:
Observations, Theory and Interpretation. Elsevier North Hol-
land, Amsterdam, pp. 436–455.
Neuberg, J., Luckett, R., Ripepe, M., Braun, T., 1994. Highlights
from a seismic broadband array on Stromboli volcano. Geophys.
Res. Lett. 21, 749–752.
Pham, V.N., Boyer, D., Kieffer, G., 1982. Sur l’existence d’une
nappe phreatique sur le flanc Sud de l’Etna au niveau du
Refuge Sapienza, et sa relation avec la structure interne du
volcan et avec la formation de la Valle del Bove. 9th
R.A.S.T., Paris, March 17–19, 1982, abstracts volume, p.
506.
Privitera, E., Sgroi, T., Gresta, S., 2003. Statistical analysis of in-
termittent volcanic tremor associated with the September 1989
summit explosive eruptions at Mt. Etna, Sicily. J. Volcanol.
Geotherm. Res. 120, 235–247.
Research Staff of INGV, Sezione di Catania, 2001. Multidisciplin-
S. Gresta et al. / Journal of Volcanology and Geothermal Research 137 (2004) 219–230230
ary approach yields insight into Mt. Etna eruption. EOS Trans.
AGU 82 (52), 653–656.
Ripepe, M., Braun, T., 1994. Air-wave phases in strombolian ex-
plosion-quakes seismograms: a possible indicator for the magma
level? Acta Vulcanol. 5, 201–206.
Ripepe, M., Gordeev, E., 1999. Gas bubble dynamics model for
shallow volcanic tremor at Stromboli. J. Geophys. Res. 104,
10639–10654.
Ripepe, M., Marchetti, E., 2002. Array tracking of infrasonic sour-
ces at Stromboli volcano. Geophys. Res. Lett. 29, 2076.
(doi:10.1029/2002g1015452)
Ripepe, M., Poggi, P., Braun, T., Gordeev, E., 1996. Infrasonic
waves and volcanic tremor at Stromboli. Geophys. Res. Lett.
23, 181–184.
Ripepe, M., Coltelli, M., Privitera, E., Gresta, S., Moretti, M., Pic-
cinini, D., 2001a. Seismic and infrasonic evidences for an im-
pulsive source of the shallow volcanic tremor at Mt. Etna, Italy.
Geophys. Res. Lett. 28, 1071–1074.
Ripepe, M., Ciliberto, S., Della Schiava, M., 2001b. Time cons-
traint for modelling source dynamics of volcanic explosions at
Stromboli. J. Geophys. Res. 106, 8713–8727.
Ripepe, M., Harris, A.J.L., Carniel, R., 2002. Thermal, seismic and
infrasonic evidences of variable degassing rates at Stromboli
volcano. J. Volcanol. Geotherm. Res. 118, 285–297.
Riuscetti, M., Schick, R., Seidl, D., 1977. Spectral parameters of
volcanic tremors at Etna. J. Volcanol. Geotherm. Res. 2,
289–298.
Rowe, C.A., Aster, R.C., Kyle, P.R., Dibble, R.R., Schlue, J.W.,
2000. Seismic and acoustic observations at Mount Erebus Vol-
cano, Ross Island, Antarctica, 1994–1998. J. Volcanol. Geo-
therm. Res. 101, 105–128.
Schick, R., Riuscetti, M., 1973. An analysis of volcanic tremor at
South-Italian volcanoes. Z. Geophys. 39, 262–274.
Schick, R., Cosentino, M., Lombardo, G., Patane, G., 1982. Volca-
nic tremor at Mt. Etna. A brief description. Mem. Soc. Geol.
Ital. 23, 191–196.
Seidl, D., Schick, R., Riuscetti, M., 1981. Volcanic tremors at Etna:
a model for hydraulic origin. Bull. Volcanol. 44, 43–56.
Taddeucci, J., Pompilio, M., Scarlato, P., 2002. Monitoring the
explosive activity of the July–August 2001 eruption of Mt.
Etna (Italy) by ash characterization. Geophys. Res. Lett. 29,
1029–1032.
Uhira, K., Takeo, M., 1994. The source of explosive eruptions of
Sakurajima volcano, Japan. J. Geophys. Res. 99, 17775–17789.
Vergniolle, S., Brandeis, G., 1994. Origin of the sound generated by
Strombolian explosions. Geophys. Res. Lett. 21, 1959–1962.
Vergniolle, S., Brandeis, G., 1996. Strombolian explosions: 1. A
large bubble breaking at the surface of a lava column as a source
of sound. J. Geophys. Res. 101, 20433–20447.
Wallis, G.B., 1969. One Dimensional Two-Phase Flows. McGraw
Hill, New York 408 pp.
Weill, A., Brandeis, G., Vergniolle, S., Baudin, F., Bilbille, J.,
Fevre, J.-F., Pirron, B., Hill, X., 1992. Acoustic sounder meas-
urements of the vertical velocity of volcanic jets at Stromboli
volcano. Geophys. Res. Lett. 19, 2357–2360.