shallow seismic attenuation and shear-wave splitting in ...hera.ugr.es/doi/15074699.pdf · ered to...

Shallow seismic attenuation and shear-wave splitting in theshort period range of Deception Island volcano (Antarctica)

Carmen Mart|¤nez-Are¤valo a, Francesca Bianco b, Jesu¤s M. Iba¤n‹ez a;c;�,Edoardo Del Pezzo b

a Instituto Andaluz de Geof|¤sica, Universidad de Granada, Campus de Cartuja s/n, 18071 Granada, Spainb Osservatorio Vesuviano, Istituto Nazionale di Geo¢sica e Vulcanologia, Via Diocleziano, 328, 80124 Napoli, Italy

c Departamento de Fisica Teo¤rica y del Cosmos, Universidad de Granada, 18071 Granada, Spain

Received 20 December 2001; accepted 25 June 2002

Abstract

The occurrence of a seismic series in Deception Island volcano (Antarctica), composed of hundreds of localvolcano^tectonic earthquakes, has permitted us to study the seismic attenuation of such a volcanic environment in theshort-distance and high-frequency range. This study has been performed using P-waves, S-waves and coda-waves andapplying different, frequency dependent and independent, techniques. The methods used for this analysis have been:spectral and broadening-of-the-pulse, for direct P- and S-waves, coda normalization for S-waves, and single back-scattering model for coda-waves. The results show that, in general, Q values are significantly smaller for the entirefrequency range used (6^30 Hz) than those found in other volcanic and tectonic areas. The attenuation for P-waves isgreater than for S-waves in the frequency independent methods, with a QL /QP ratio that ranges between 1.9 and 3.2.Comparing the Q-factor obtained for S-waves we have observed clear differences as a function of the method used;the coda normalization method has supplied significantly higher Q values (Qd ) than the other two methods (QL ). Wehave interpreted this discrepancy as an effect of the methods: coda normalization and single back-scattering methodseliminate the contribution of the near-surface attenuation in their Q values. Comparing both QL and Qd we haveestimated the near-surface attenuation under the recording site, named QU . On the other hand, we have observed thatQd has anomalous frequency dependence, with a minimum value at 21 Hz. This pattern is interpreted as an effect ofstrong scattering of the seismic waves in the source area of the earthquakes. Qc values depend clearly with frequencyand lapse time and the lapse time dependence could be interpreted as a depth dependence of the seismic attenuation inDeception Island volcano, Antarctica. The obtained Q values have allowed us to separate the contribution of intrinsicand scattering attenuation, deriving that the scattering attenuation is predominant over the intrinsic effects. Finally, inorder to investigate how the heterogeneous medium of the volcanic island could produce other effects, we havechecked whether it produces polarization of the shear-waves. The preliminary results of the polarization directionindicate a main E^W strain direction. All these evidences reveal the strongly heterogeneous structure of DeceptionIsland volcano.< 2003 Elsevier B.V. All rights reserved.

Keywords: Seismic attenuation; shear-wave splitting; volcanic seismicity; short period range; Antarctica

0377-0273 / 03 / $ ^ see front matter < 2003 Elsevier B.V. All rights reserved.doi:10.1016/S0377-0273(03)00248-8

* Corresponding author. Fax: +34-958160907. E-mail addresses: [email protected] (C. Mart|¤nez-Are¤valo),[email protected] (F. Bianco), [email protected] (J.M. Iba¤n‹ez), [email protected] (E. Del Pezzo).

VOLGEO 2685 4-11-03 Cyaan Magenta Geel Zwart

Journal of Volcanology and Geothermal Research 128 (2003) 89^113

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1. Introduction

Geological structures characterizing volcanicenvironments are strongly heterogeneous; theyare composed of a complex distribution of ele-ments of di¡erent sizes and properties. We canobserve the presence of lava £ows, low velocityash layers, complex fracture patterns withoutmain directions, etc. At the surface, bombs,coarse scoria or viscose lavas can be found nearthe emission centers, whereas ¢ne ash of smallmaterials is to be found over wide areas. Atdepth, intrusion of magma bodies, hydrothermaldeposits and anomalies, among others, draw acomplex structure picture. This complex structurea¡ects the seismic waves propagation. Almendroset al. (2002) and Saccorotti et al. (2001), demon-strated that rough topography and a laterallyvarying shallow structure distorted the signal,producing phenomena like focusing and defocus-ing of the seismic rays, wave scattering, amplitudeabsorption, and others. All these phenomena areparameterized by the Q-parameter (or quality fac-tor) which accounts overall for the energy decayas a function of distance. The Q-factor assumesdi¡erent values, depending on the wave phasesused in the analysis, and/or by the geologicalcharacteristics of the area under study. In general,Q values are signi¢cantly lower for volcanic envi-ronments than for tectonic areas. Further, Q de-pends on frequency in a di¡erent manner for vol-canic regions than for tectonic areas. Thisdependence is also di¡erent according to the dis-sipation mechanisms accounted for. Main mecha-nisms generally considered are intrinsic dissipa-tion and seismic scattering, for which twodi¡erent types of Q-factor can be de¢ned: Qi

(the intrinsic Q) and Qs (the scattering Q). Maye-da et al. (1992) observed an anomalous behaviorof Qi and Qs for Hawaii in comparison to CentralCalifornia or Long Valley Caldera. In some vol-canic areas, at shallow depths, e.g. Mt. Etna orCampi Flegrei (Del Pezzo et al., 1995), scatteringattenuation seems to prevail over intrinsic dissipa-tion. At Deception Island volcano Saccorotti etal. (2001) observed strong deviations of thewave front due to a velocity contrast in the upperlayers of the volcano. Vila et al. (1995) measured

low Qc values (Qc = 10.6Wf1:32 in the frequencyrange between 1 and 14 Hz).

The aim of the present work is the experimentaldetermination of the Q-factors at Deception Is-land in the high frequency range (up to 30 Hz),and short hypocentral range, using P-waves,S-waves and coda-waves. The study of the directP- and S-waves attenuation is motivated by:(1) there are no attenuation studies in the islandusing these waves; (2) we try to compare the P-and S-waves seismic attenuation, using the sametechniques, and to interpret the results in terms ofthe volcanic structure; (3) we would like to testthe stability and accuracy of techniques and datain order to use them in a future attenuative to-mography of the island. For coda waves the studyis focused on the comparison of the results withthose obtained using direct S-waves, and withthose derived previously by Vila et al. (1995).Then, we separate the intrinsic and scattering ef-fects on the seismic attenuation and investigatethe e¡ect of the near-surface attenuation overthe di¡erent Q values derived previously. Theabove analysis shows that the attenuative struc-ture of Deception Island is complex, with anom-alous frequency dependence, and strong di¡eren-ces of the Q values according the used method.To check the possible relationship between com-plex Q-patterns and structure we perform prelimi-nary observations of shear-wave splitting.

This work has been structured as follows: Firstwe describe the data set used and its spectral char-acteristics. Then we summarize the method used,and ¢nally the results are discussed in the light ofthe separate estimates of Qi and Qs. Surface het-erogeneities, evidenced by the observed splittingphenomena, will be shown to be the main causeof the observed prevalence of scattering attenua-tion.

2. Deception Island volcano and seismic datadescription

Deception Island is the most active and impor-tant volcano of the South Shetland Islands (Ant-arctica), and it is located at 62‡59PS and 60‡41PW,northeast of the Antarctic Peninsula. It is consid-


C. Mart|¤nez-Are¤valo et al. / Journal of Volcanology and Geothermal Research 128 (2003) 89^11390

ered to be the main active volcano of the back-arcbasin of the Brans¢eld Strait (e.g. Baker, 1990).This caldera island (see Fig. 1) produced severaleruptions over the last 200 years. During the1967^1970 period three eruptions took place de-stroying the Chilean and British bases. Since 1994the seismic activity of the island has been moni-tored, using one or more seismic arrays andsparse seismic networks, under the framework of

the Spanish Antarctic Program. Alguacil et al.(1999), Almendros et al. (1997, 1999) and Iba¤n‹ezet al. (1997, 2000) provided a dense and exhaus-tive description of the seismicity recorded duringthe 1988^1999 period.

Data analyzed in the present work belong to aseismic swarm that occurred in January^February(1999) when more than 1500 volcano^tectonicearthquakes were detected. This swarm has been

Fig. 1. (a) Map of the South Shetland Islands region, Antarctica, showing the position of Deception Island. (b) Map of Decep-tion Island, showing the location and con¢guration of the seismic antennas used during the 1998^1999 experiment (c).


C. Mart|¤nez-Are¤valo et al. / Journal of Volcanology and Geothermal Research 128 (2003) 89^113 91

analyzed by Iba¤n‹ez et al. (2003). Seismicity wasrecorded by a dense seismic array composed of 10vertical and three 3-component short-period seis-mometers with natural period at 4.5 Hz, and with£at response electronically enlarged between 1^50Hz (see Iba¤n‹ez et al., 2003; Saccorotti et al., 2001)for a more detailed description of the site andinstruments. Location of these events has beenperformed by Iba¤n‹ez et al. (2003). In Fig. 2 thehypocentral map of the analyzed events is re-ported. Digital records were sampled at 200 sam-ples/s, allowing us to study the attenuation inprinciple up to 50 Hz. The short hypocentral dis-tance, on average lower than 5 km, and the lowmagnitude of the events, between 30.5 and 2.0,therefore, the high values of corner frequencies,allows us to this high frequency analysis.

Prior to the attenuation study we analyzed thesignal to noise ratio to delineate the frequencyrange of the study. We compared the pre-event

spectral amplitude to the P-waves, S-waves andcoda-waves spectral amplitudes, for a set of se-lected events at di¡erent distances and magni-tudes. In Fig. 3 we show two examples for twoearthquakes, the ¢rst one with Mw = 1.0 and thesecond with Mw = 0.0, and hypocentral distancesmaller than 4 km for both of them. The compar-ison shows that the most stable frequency rangefor the analysis corresponds to a band between 6and 30 Hz. Below 6 Hz the contribution of thebackground noise is so high that the spectral am-plitude of the pre-event noise is comparable to thespectral amplitude of an earthquake of magnitude1.0. The source of this low frequency noise is thecombined e¡ect of oceanic noise and volcanictremor (Iba¤n‹ez et al., 2000). Over 30 Hz the signalto noise ratio value also becomes insu⁄cient,therefore our analysis, about the seismic attenua-tion of Deception Island, will be centered in the6^30 Hz frequency band.

Fig. 2. Map of the earthquake locations analyzed in the present study, obtained by Iba¤n‹ez et al. (2003). The position of the seis-mic array is shown in the ¢gure.



3. Description of the methods

In the present section we shall describe shortlythe methods used to estimate the seismic attenu-ation. We apply the broadening-pulse and thespectral decay methods for both P- and S-waves,the coda normalization method for S-waves, andthe single scattering model for coda-waves.

3.1. Broadening-of-the-¢rst-pulse method

The measurement of direct wave pulse durationo¡ers us information about the source and themedium that the waves crossed from their source

to the receiver. The method has the advantagethat it is not necessary to know the instrumentalresponse and precise measurement of pulse ampli-tudes. The method only requires a half length ofthe ¢rst direct waves (P- or S-waves). Gladwinand Stacey (1974) applied, for P-waves, an empir-ical relation that relates the pulse duration, thesource and the travel path of a seismic wave:

d 1=2 ¼ d 0 þctQP

¼ d 0 þcdvPQP

ð1Þ

where d1=2 is the P-pulse duration at a distance dfrom the source, d0 is the P-pulse duration at thesource, c is a constant, t is the travel time, QP is

Fig. 3. Examples of two local earthquakes recorded at the Fumaroles array, comparing the signal to noise ratio for P-waves,S-waves and coda-waves. (a) Seismogram of an earthquake of magnitude 1.0. (b) Earthquake of magnitude 0.0. Both events werelocated at a hypocentral distance smaller than 4.0 km. Over each seismogram three blocks show the window used to estimate thespectra, noise, P-waves, S-waves and coda-waves.



the quality factor for P wave, and vP is the P-wavevelocity. The pulse duration is de¢ned as the lin-ear extrapolation of the maximum slope to thetime axis and the ¢rst zero (Liu et al., 1994).Eq. 1 is valid for homogeneous media, but Wuand Lees (1996) have probed that this equationcan be applied to fractured media, as in a volcanicenvironment. Jongmans (1991) found that thismethod is not valid when the hypocentral distan-ces are less than 1.2 times the wave length. Of theexperiments with ultrasonic acoustic pulses, in thekHz frequency band, Gladwin and Stacey (1974)determined the value of the constant c to be equalto 0.5.

To apply the broadening wave pulse method wehave to observe that there is not a relationshipbetween the pulse duration and the magnitude.Consequently, we can rule out that the pulsebroadening is due to the source e¡ects. Secondly,the earthquake sources have to be of the sametype, and therefore the pulse form and widthmust be the same for all earthquakes at thesource.

3.2. Spectral decay method

The method assumes that direct wave displace-ment spectra (P- or S-waves) are not £at, for fre-quencies below the corner frequency, due to thecontribution of attenuation. We can divide themedium into two layers, the ¢rst being the near-surface layer and the other including the rest ofthe travel path of the seismic waves, each onewith a constant quality factor, Q. It is possibleto express the spectral amplitude as:

Aðf ; tshallow; tdepthÞ ¼ A0ðf Þe3Z ftshallow

Qshallowþ3Z ftdepth

Qdepth ð2Þ

where Q31shallow and Q31

depth are the attenuation inthe shallow and deep layers, respectively, tshallow

and tdepth are the travel time in each layers andA0(f) is the source spectral amplitude. We de¢neQtotal as:

tQtotal

¼ tshallow

Qshallowþ tdepth

Qdepthð3Þ

where t is the total travel time. The term tshallow/

Qshallow is commonly named U in studies of localseismicity (Anderson and Hough, 1984; Aber-crombie, 1997; Margaris and Boore, 1998) andthe t* name is used only for teleseismic distances.Therefore, we shall rename in the present workQshallow as QU . If we substitute Eq. 3 in Eq. 2and we take natural logarithms:

lnðAðf ; tÞÞ ¼ lnðA0Þ3Z ft=Qtotal ð4Þ

we obtain the equation of a straight line where theindependent variable is the frequency, the depen-dent variable is the ln(A(f,t)), and the slope is theterm 3Zt/Qtotal, which is independent of the fre-quency. The term Qtotal engulfs the contributionof the near-surface and deepest attenuation. If weare able to estimate the attenuation in the deepestlayer by another method, we can calculate thenear-surface attenuation, as Havskov et al.(2002) have recently proposed.

3.3. Coda normalization method

The S-wave attenuation can be estimated fol-lowing Aki (1980). The method is based on theidea that at a lapse time much greater than theS-wave travel time the seismic energy is uniformlydistributed in a volume surrounding the source.The limits of this assumption, of empirical origin,have been investigated theoretically in the frame-work of multiple scattering processes based on theradiative transfer theory (Sato and Fehler, 1998).

Interpreting S-coda as a random superpositionof scattered S-waves (Aki, 1980), the S-coda en-ergy as a function of the source, site and pathe¡ects, can be written as:

ACoda Sðf ; tÞOWSi ðf ÞMNS

j ðf ÞMe3Q

31c Z f t

tnð5Þ

where ACoda S(f,t) is the S-coda spectral amplitudeat the frequency f, Wi

S(f) the source i energy ra-diation at the same frequency f, MNS

j (f)M is theS-wave site ampli¢cation factor for site j, n takesinto account the geometrical spreading factor,and takes di¡erent values depending on the dom-inance of surface, body and di¡usive waves (be-tween 0.5^1.5, respectively), and e3Q

31c Z f t is the

attenuation function, which accounts for the



propagation process where Qc31 is the S-coda

wave attenuation.The S-wave spectrum for the ith source and jth

station (at the distance rij ) can be written as:

AS Directðf ; tÞOWSi ðf Þr2ij

MNSj ðf ÞMeð3Q

31d Z f rijÞ=ðv0Þ ð6Þ

where Qd31 is the direct S-wave attenuation, v0 is

the S-wave velocity and f is the frequency. Eval-uating Eq. 5 at ¢xed lapse time tc we can write :

lnrijMAS Direct ðf ; tÞMACoda S ðf ; tcÞ

¼ 3ðQ31d ðf ÞZ f =v0Þrij þ cte ð7Þ

A least square regression analysis of the left-hand side of the relationship (Eq. 7) vs. hypocen-tral distance allows us to estimate Qd

31 from theslope. Using this method, the site and source ef-fects and S-coda attenuation are cancelled by theratio (Eq. 7). For more details, see Aki (1982).

The spectral amplitude of the direct S-wave andcoda-wave can be estimated either evaluating theFourier transform of the signal in a windowaround the direct S-wave arrival time, by the di-rect S-wave, and the ¢xed lapse time, by the coda-wave, or, ¢ltering the signal with a pass-band ¢l-ter and calculating the rms of the signal amplitudein the window centered at direct S-wave arrivaltime, by direct S-wave, and at lapse time tc, bythe coda wave.

3.4. Single back-scattering method for coda waves

In the assumption of a single scattering process,and for a source placed in the same position asthe receiver, the coda envelope, as a function ofthe time t elapsed from the origin time (namedlapse time), can be expressed, following (Akiand Chouet, 1975), by:

Aðf ; tÞ ¼ A0ðf Þt3ne3Z ft=Qc ð8Þ

where A0(f) is a term that depends on the source,travel path and site geology, t3n is the geometricalspreading function, and Qc is the coda-wave qual-ity factor. Qc

31 is estimated by taking the loga-rithms of both sites of Eq. 8:

lnðAðf ; tÞtnÞ ¼ lnðA0ðf ÞÞ3Z ft=Qc ð9Þ

and then least squares ¢tting the measured enve-

lope. The slope is 3Zt/Qc and the independentterm is ln(A0(f)). Qc can be alternatively deter-mined by a non-linear ¢t of the relationship ofEq. 8 directly (see e.g. Iba¤n‹ez et al., 1993).

4. Application of the methods and techniques

The ¢rst step to estimate the seismic attenua-tion was the selection of the seismic stations, be-cause, as mentioned, the earthquakes were re-corded by a seismic array. Since the aperture ofthe array was not longer than 300 m, at least anorder of magnitude less than the hypocentral dis-tance, we should expect similar results among allthe stations. We have selected the station 6G asthe reference station for vertical displacements,and in case of using horizontal components F1-3D was the station selected (see Fig. 1).

4.1. Broadening-of-the-pulse method

4.1.1. P-wavesAs described above, the technique is not valid

for hypocentral distances shorter than 1.2 timesthe wave length. We checked the typical wavelength of the ¢rst P-wave pulse and found thatthe average duration of the ¢rst cycle is 0.08 s,which correspond to a frequency of 12.5 Hz. As-suming an average P-wave velocity of 3.7 km/s,our typical wave length is 300 m, and thereforedistances shorter than 366 m cannot be used. Inorder to be certain that at all time we shall be atdistances greater than 1.2 times the wave length,we shall start our analysis at 1 km of hypocentraldistance. To satisfy the condition of the samepulse duration at the source we must restrict ourdata set in a uniform range of magnitude. Weselected a magnitude interval between 1.0 and2.7, in which the pulse duration appeared clearly.The data set analyzed was composed of 140 earth-quakes distributed in a hypocentral range between0.5 and 8.5 km.

In the literature there are several ways to esti-mate the pulse duration. In the present study weused the Liu et al. (1994) de¢nition in which thepulse duration is de¢ned as the linear extrapola-tion of the maximum slope of the beginning of the



Fig. 5. Pulse duration for P-waves against magnitude (a) and polarity of the ¢rst motion direction (b). As observed, there is no aclear relationship between the duration and the other parameters.

Fig. 4. An example of the procedure followed to estimate the P-wave pulse duration. (a) A 0.25-s window of the arrival of theP-wave. (b) The half pulse duration and the measurement procedure. d indicates the pulse duration considered.



pulse before the ¢rst onset, using as reference thebaseline, and the ¢rst zero after the maximum ofthe pulse (see Fig. 4). This measure is done on thevelocity seismogram, therefore this value shouldcorrespond to a half pulse of the displacementseismogram. In order to check the possible depen-dence of the pulse duration on the magnitude ofthe events we have plotted the pulse durationagainst hypocentral distance as a function of themagnitude. As observed in Fig. 5a, there is nocorrelation between the pulse duration and mag-nitude. Iba¤n‹ez et al. (2003) reveals the complexityof the source mechanism for the seismic sequenceanalyzed in the present study. One of the eviden-ces of this complexity is the presence of di¡erentpulse directions (up or down) for earthquakes in avery short distance interval from the array. Tocheck if this complexity could a¡ect our results,we plotted the pulse duration against the hypo-central distance as a function of the polarity ofthe signal (see also Fig. 5b), showing that there isno relationship between ¢rst pulse direction andduration of the pulse.

We estimated QP ¢tting the experimental datato the relationship of Eq. 1 (see Fig. 6). The cor-

relation coe⁄cient obtained was enough to con-sider the result as signi¢cant, b= 0.84. Then, weestimated QP, ¢xing the value of c and vp. For ourstudy we shall use the c value of 0.5 because thisvalue has been experimentally and numericallytested. For example, Gladwin and Stacey (1974),analyzing measures on acoustic pulses propagat-ing in massive rocks, obtained a c value of 0.53 Q0.04. Kjartansson (1979) presented a linear modelof rock inelasticity characterized by a Q exactlyindependent of frequency. From his equation it ispossible to derive a c value of 0.485. The P-wavevelocity used in this study has been obtained as anaverage value, for the depth and distance rangeused derive from the velocity model presented byIba¤n‹ez et al. (2003) for Deception Island volcano.Using vp = 3.7 Q 0.1 km/s and c= 0.5 the derivedQP result is :

QP ¼ 28 � 9

4.1.2. S-wavesThe S-wave pulse was measured in the radial

projection of the horizontal components of sta-tion F1-3D (see Fig. 1). Then, we selected the

Fig. 6. Plot of the pulse duration for P-waves vs. distance, showing the best ¢t and the QP value.



earthquakes with an incidence angle lower than37‡. The purpose was to have as pure as possibleS-wave incidence without the in£uence of anyconverted waves. The value of the angle hasbeen obtained observing the ratio between P- andS-waves velocity (vp = 3.7 Q 0.1 km/s, and vs =2.24 Q 0.07 km/s) and the condition of critical an-gle. With these strict conditions we reduced theinitial data set from 863 events to 80, but only23 of them have clearly visible S-wave ¢rst pulses.In Fig. 7 we show an example of the S-wave overthe radial component and the zoom of the pulse.The method used to calculate the pulse durationwas the same used for P-waves. As in the case ofthe P-waves pulse, we tested the possible depen-

dence of the pulse duration with magnitude orfocal mechanism ¢nding that there is not any re-lationship. The ¢tting of the pulse durationagainst hypocentral distance is reported in Fig.7. The correlation coe⁄cient obtained for this ¢tis b= 0.64, which implies to consider the result asstatistically signi¢cant. The QL value has to beobtained using a c constant and the shear-wavevelocity, vs. There is no reported informationabout the constant c in the literature. To checkthe in£uence of c over QL we used di¡erent val-ues, from 0.5 to 3, as reported in Table 1. Theway to decide the best c value of shear-waves iscomparison of the Q values, derived by thepresent method, with those obtained using other

Fig. 7. An example of the procedure followed to estimate the S-waves broadening of the pulse. (a) An example of a S-wave pack-age and the half-pulse duration. (b) The half pulse duration and the measurement procedure. (c) Plot of the pulse duration forS-waves vs. distance, showing the best ¢t and the QL value.



independent techniques such as the spectral decaymethod, discussed below.

4.2. Spectral decay method

4.2.1. P-wavesThe majority of the 860 located earthquakes

show a clear P-wave package, before the arrivalof the S-waves, with a good signal to noise ratio.This observation has permitted us to apply thespectral decay method for the majority of the

events, using the SEISAN software package (Hav-skov and Ottemo«ller, 1999). The procedure of thecalculation in station 6G was the following.Firstly, a band pass ¢lter, between 6 and 15 Hz,was applied to all data. This condition, at lowfrequency, was established in order to minimizethe in£uence of the background seismic noise.At high frequency, we cut the signal below theaverage corner frequency, located around 20 Hz(Havskov et al., 2003; Iba¤n‹ez et al., 2003). Sec-ondly, the P-wave window had a ¢xed duration of0.52 s (Fig. 8). This condition implies that allearthquakes with S^P time smaller than this valuewere rejected. Thirdly, in all cases, we rejectedthose spectra that, inside of the 6^15 Hz interval,show a signal to noise ratio lower than 2.Fourthly, all spectra were visually inspected andthose that did not show a continuous decay wererejeceted, as some spectra with a positive slopewere observed. Probably, their signal containedthe arrival of other phases. In total, 523 earth-quakes were analyzed. After the calculation of

Fig. 8. Two examples of the application of the spectral method to estimate Q using P-waves (a) and S-waves (b). Over each seis-mogram the window used in the spectral analysis is marked. In each spectrum we marked two lines. The horizontal one repre-sents the expected £at spectrum in case of no attenuation, and the other one is the estimated ¢t used to calculate Q. From thedi¡erence between both slopes spectral-Q is calculated. The decay of the spectrum after the observed corner frequency is repre-sented as well.

Table 1QL as a function of the constant c

c QL vQL

0.5 26 Q 91.0 52 Q 181.5 80 Q 302.0 100 Q 402.5 130 Q 503.0 160 Q 50

vQL is the standard deviation of the estimated QL values.



the slope of the spectral decay, those results thatprovided a correlation coe⁄cient greater than 0.9were selected, reducing the initial data set to a¢nal number of 200 earthquakes. The obtainedU value from this analysis was 0.05 Q 0.02, thatfor the average hypocentral distances and the ve-locity model used provides a QP value of QP =13 Q 6.

4.2.2. S-wavesIn order to test if there are any di¡erences be-

tween the analysis using vertical or horizontalcomponents we performed the study in the verti-cal 6G station, and in the two horizontal compo-nents of the F1-3D station. Following the samecriteria described for P-waves application, theU value for vertical displacement, using 147 data,was found to be 0.04 Q 0.02. For the N^S andE^W components, U was 0.05 Q 0.02. For the hy-pocentral range of the selected data, the QL val-ues are: QL (station 6G) = 41 Q 19, and QL (hor-izontal components) = 35 Q 13.

As observed, both values are located inside theerror interval and can be considered as the same.

Comparing the present results to those describedin Table 1, we can estimate the best c value of theS-wave broadening-pulse method. A c value of1 provides the better agreement between bothmethods.

4.3. Coda normalization method

The application of this method requires the se-lection of two windows over the seismogram; the¢rst one centered over the S-waves, and the sec-ond over the coda in a reference time, to calculateAS Direct(g) and ACoda S(g) of the relationship inEq. 7. This analysis was done in the frequencydomain in 12 frequency bands, starting at 7 Hzand ¢nishing at 29 Hz, with a half window length,vf, of 1 Hz. The ¢ltering was done using an 8-poleButterworth ¢lter. At every frequency band,AS Direct(g) and ACoda S(g) were estimated as theRMS of the amplitude of the ¢ltered trace, fol-lowing the work of Del Pezzo et al. (1995), with awindow length of 2 s for both of them. The refer-ence lapse time for the coda window was ¢xed at8 s (see Fig. 9), due to the small size of the ana-

Fig. 9. An example of the ¢t of the coda normalization method, for 21 Hz, vs. the distance. (a) In the seismogram the referencecoda used in the analysis has been marked. (b) Small dots represent the whole data set, and darker dots represent the averagevalues obtained using a moving window over the raw data.



lyzed events, which implies a short coda duration.The selection of the data set was done as follows:(1) a signal to noise ratio, at all the frequencybands, greater than 2 for the S and coda windows,and (2) the coda shape has to decay uniformly.With these criteria we selected a subset of 400earthquakes. The conditions reduced the hypocen-tral range from the initial [0.5^10 km] to [0.5^3.0km]. Once the ratio AS Direct(g)/ACoda S(g) wasperformed at the di¡erent frequency bands, we

¢tted this ratio vs. hypocentral distance. The ¢twas done using two procedures: (1) ¢tting thewhole data set as a function of the distance, and(2) transforming the data into a new set using anaverage moving window (see Fig. 9) of 0.3 kmlong, and averaging all the ratios AS Direct(g)/ACoda S(g) ; this window was moved at 2/3 of itslength. Both ¢ts, with raw data and using themoving window, supply similar results, but thecorrelation coe⁄cient is sensibly improved usingthe moving window method. The obtained Q-fac-tor, Qd , as a function of the frequency is shown inFig. 10 and Table 2. Although the initial fre-quency interval analyzed was between 6 and 30Hz, the three ¢rst bands (6^8, 8^10 and 10^12Hz), and the last one (28^30 Hz) show a verylow correlation coe⁄cient, less than 0.4, andthey were not taken into account. The frequencydependence of Qd cannot be represented by a sim-ple linear law, because two patterns are visible:between 13 and 21 Hz, Qd decreases slowly withthe frequency, and for higher frequencies Qd sgrows. The Qd values are signi¢cantly greaterthan the QL ones derived from the spectral orthe broadening-of-the-pulse methods.

Table 2Qd values using the coda normalization method

Frequency Qd vQd Correlationcoe⁄cient, b

(Hz)

13 140 Q 50 0.6615 140 Q 40 0.7417 120 Q 30 0.8219 98 Q 16 0.9021 93 Q 12 0.9423 133 Q 21 0.9025 190 Q 40 0.7827 300 Q 110 0.55

vQd is the standard deviation of the estimated Qd values.

Fig. 10. Frequency dependence of Qd , Q of S-waves estimated using the coda normalization method. Error bars represent the es-timated error obtained by applying the error theory over the relationship in Eq. 7.



4.4. Single back-scattering method

We selected those earthquakes which their codawaves presented a uniform decay in a long inter-val, without the interference of other arrivals, ob-taining a ¢nal set of 500 earthquakes. The analy-sis was done as a function of the frequency andlapse time. To compare these results to the codanormalization analysis, we used the same fre-quency bands, between 6 and 30 Hz. The lapsetime dependence was calculated from 6 to 22 s, inintervals of 2 s. In the present work we use theterm lapse time as the time interval between theorigin time of the earthquake and the time of theend of the analysis. The start time of every coda

was ¢xed using the criteria de¢ned by Iba¤n‹ez etal. (1993): the start time is the point in which theterm ln (ACoda(g,t)Wt) starts to decrease afterreaching its maximum value. The end of thecoda corresponds to the window ¢xed by thelapse time studied if the ratio of the RMS ofthe signal and the average pre-event noise wasgreater than 1. For each lapse time the coda en-velopes were estimated, after ¢ltering the traces,using a 0.5-s long moving window, and then itwas moved 50% of its length along the coda;at every window the RMS of the seismogramwas calculated. Qc value was obtained ¢ttingln(ACoda(g,t)Wt) against t, and using the non-linearmethod of Iba¤n‹ez et al. (1993). Both analyses pro-

Fig. 11. An example of the coda envelope for the seismogram represented in (a). Darker area shows the portion of the coda ana-lyzed. In the plot (b) we compare the log^log and the non-linear ¢t observing that for this case both methods provide similar re-sults.



vided similar results when the signal to noise ra-tio, at the end of the coda, was two or more (Fig.11). Otherwise, the linear method provides sys-tematically overestimation of Qc, as described byIba¤n‹ez et al. (1993). We selected those Qc valuesthat simultaneously provided the correlation coef-¢cient, for the linear and non-linear methods,greater than 0.9, obtaining similar Qc values, in-dependent of the method. At lapse times of 6 and22 s the number of data analyzed was not enoughto provide statistical signi¢cance. In Fig. 12 we

show the variation of Qc with the frequency forthe di¡erent lapse times, and in Fig. 13 we plotthe dependence of Qc with the lapse time for thedi¡erent frequency bands, observing a regulargrowing of Qc. It is possible to obtain a relation-ship of Qc with frequency following the law:

Qc ¼ Q0 ðf =f 0Þn

where Q0 is the Qc value at the frequency of 1 Hzand f0 is the reference frequency of 1 Hz. In Table3 we show the results for the di¡erent lapse times.

Table 3Qc laws as a function of the frequency

Lapse time Q0 QvQ0 nQvn Correlation coe⁄cient, b(s)

8 12.7 Q 0.3 0.9 Q 0.2 0.9110 24.6 Q 0.2 0.70 Q 0.04 0.9812 30.1 Q 0.2 0.66 Q 0.03 0.9914 27.0 Q 0.2 0.73 Q 0.03 0.9916 26.5 Q 0.2 0.77 Q 0.04 0.9818 26.1 Q 0.3 0.79 Q 0.05 0.9720 20.0 Q 0.4 0.9 Q 0.1 0.92

vQ0 is the standard deviation of the estimated Q0 values.vn is the standard deviation of the estimated n values

Fig. 12. Variation of Qc with the frequency for all the lapse time used in the present study, from 8 to 20 s.



5. Discussion

In the present work we have analyzed the shortperiod seismic attenuation of Deception Islandvolcano, using data recorded in a seismic antenna,belonging to a seismic series occurring in Janu-ary^February 1999. The Q value has been ob-tained using P-waves, S-waves and coda-waves.

5.1. Direct waves

The QP values (Table 4), derived in the presentwork, are low, as expected for volcanic areas

where high fracturation and the nature of theupper layers should produce a high attenuation.Although both values seem to be slightly di¡erent,it does not a¡ect the interpretation of the results,a high P-wave attenuation. These QP values areindependent of the frequency, but they have asimilar frequency range of application: the spec-tral method was applied in the 6^15-Hz band, andthe duration of the analyzed half pulse of P-wavesranges between 0.020 and 0.075 s, which corre-sponds to frequencies between 6 and 25 Hz. Incomparison to other volcanic areas, our resultsare similar to or slightly smaller than those which

Table 4Q values, for both P- and S-waves, obtained with the frequency independent methods

Method QP vQP Q31P QL vQL QL

31

Broadening 28 Q 9 0.036 52* Q 18 0.019Spectral 13 Q 6 0.077 41 Q 19 0.024

This value has been selected using the constant c= 1.vQP is the standard deviation of the estimated QP values.vQL is the standard deviation of the estimated QL values.

Fig. 13. Qc lapse time dependence for di¡erent frequency bands, from 7 to 29 Hz.



can be found in the literature; for example, forthe Vesuvius volcano Bianco et al. (1999), usingsimilar distance ranges, obtained a QP value of 42for the broadening-of-the-pulse method and QP of36 for the spectral method. QL (Fig. 14 and Table4) is comparable to the work of Bianco et al.(1999) for the Vesuvius volcano area. They ob-tained a QL of 59, and this value is 41 Q 19 inthe present work. These results re£ect similar be-havior, in attenuation, of both volcanic areas.

The ratio QL /QP, for the spectral method, isclose to 3.2, clearly higher than the results ofBianco et al. (1999) who obtained a ratio oftwo, but comparable to the results of Yoshimotoet al. (1993) for the Kanto area in Japan. For thebroadening-of-the-pulse method, the ratio is 1.9,similar to those found by several authors in othervolcanic areas (see Sato and Fehler, 1998). Thedi¡erences between both ratios are mainly relatedto the QP values derived in the spectral method.In the P-waves spectra we observed that manyearthquakes have positive slopes of the spectralamplitude with respect to the theoretically ex-pected negative slopes, similar to the observation

by Bianco et al. (1999). In our study, we elimi-nated all these events with this bump in the spec-tra. The use of these discarded data in the estima-tion of QP could provide an increase of the QP

value, because we reduce the slope of the averagespectra, and therefore a decrease of the ratio QL /QP.

5.2. Coda normalization and coda-Q results

The calculated Qd values have an anomalousbehavior as a function of frequency. In the 11^21-Hz band Qd decreases slowly with the fre-quency, and between 21 and 29 Hz, Qd clearlygrows. The minimum Qd value, which corre-sponds to the highest attenuation, is located at21 Hz. Following the results derived by Iba¤n‹ezet al. (2002), the analyzed seismic series is mainlyclustered at a focal depth of 2^3 km, and at thisdepth the average S-waves velocity is 1.3 km/s(Saccorotti et al., 2001). With this velocity, andusing the frequency of 21 Hz, we derived that thepredominant average wave length of the analyzedseismic waves is 62 m. Iba¤n‹ez et al. (2003) derived

Fig. 14. A comparison of the di¡erent Q values estimated for S-waves using the spectral, broadening-of-the-pulse, and coda nor-malization methods. Error bars represent their estimated standard deviations.



the characteristic fault lengths of the seismic se-ries, and this length was ¢xed at 60 m. The co-incidence between both values allows us to deter-mine that the peak in attenuation, observed atthis frequency and for this method, could be pro-duced by a strong scattering e¡ect of the seismicwaves close to the source region.

The Qc values derived in the present work aredi¡erent from those obtained by Vila et al. (1995)(see Table 5 and Fig. 15). As mentioned previ-

ously, these authors found Qc = 10.6 f1:32, for the1^15 Hz frequency interval and for a lapse timeshorter than 16 s. The main di¡erence is related tothe frequency dependence, our frequency depen-dence varying between 0.79 and 0.86, with anaverage value of 0.83. In Fig. 16 we compareour results with those obtained by other authorsin other volcanic regions. As observed, Etna andVesuvius volcanoes show higher Qc values thanDeception Island volcano. These di¡erences canpossibly be explained by assuming a relationshipto the short epicentral distances used in thepresent work.Qc clearly depends on the frequency and lapse

time for Deception Island volcano (Figs. 12 and13; Table 6). This lapse time dependence can beinterpreted as a depth dependence of the seismicattenuation. The main data analyzed have a focaldepth of around 2^3 km, and only a few of themare deeper than 3^4 km. When the lapse timegrows, the volume of medium involved in the

Fig. 15. Comparison of Qc values obtained by Vila et al. (1995) with those obtained in the present work for two extreme lapsetime intervals, 8 and 18 s. Over each single Qc value we have represented the errors bars derived by their standard deviation.Straight lines represent the best ¢t of Qc with frequency.

Table 5Comparison of Qc values for Deception Island at a lapsetime of 16 s

Frequency Vila et al. (1995) This work(Hz)

7 138 129 Q 99 193 143 Q 911 251 160 Q 1013 313 180 Q 1015 378 204 Q 11



scattering process grows, in surface and in depth.For example, for the lapse time of 8 s, and usingthe average S-waves velocity of this area (1.3 km/s), the volume of medium that could contribute inthe generation of coda waves, under the assump-tion of a single back-scattering process, is an el-lipsoid of revolution with major semi-axis of 5.2km. For the larger lapse time (20 s), the semi-axishas to be 13 km long. The dimensions of Decep-

tion Island are similar to the surface projection ofthe greatest ellipsoid. We should expect thateverywhere the attenuative e¡ects of the surfacestructure of the island were similar, and thereforethe variation of Qc with lapse time should becaused by the enlargement, in depth, of the revo-lution ellipsoid. In fact, the structural works ofGrad et al. (1992, 1993) or Saccorotti et al.(2001) show a clear variation of the wave’s veloc-ity with depth, and this variation should be re-£ected in the values of seismic attenuation.

Bianco et al. (1999), in their study of the seis-mic attenuation in the Vesuvius volcano, pointedout the possible contribution of surface waves oncoda formation. Their hypothesis is based, amongothers, on the shallow hypocentral distribution ofthe analyzed seismicity, a focal depth shallowerthan 3 km. In our case, we are in a similar sit-uation, because 90% of our data have a focaldepth between 2 and 3 km. Therefore, it couldbe possible to study the variation of our resultschanging the geometrical spreading values from

Fig. 16. Comparison of the Qc values for Mt. Etna (Del Pezzo et al., 1995) and Vesuvius (Bianco et al., 1999) with the presentwork.

Table 6Evidences of the lapse time and frequency dependence of Qc

for Deception Island volcano

Lapse time Qc at 11 Hz Qc at 25 Hz(seconds)

8 ^ 197 Q 1410 146 Q 14 225 Q 1312 145 Q 11 255 Q 1514 155 Q 11 283 Q 1616 162 Q 10 312 Q 1818 168 Q 10 371 Q 2120 172 Q 9 ^



n= 1 to n= 0.5 under the hypothesis of changefrom body to surface waves. In order to checkthe possible Q di¡erences according to the geo-metrical spreading value, we calculated again Qc,for a lapse time of 10 s, as an example. The newestimations follow the law:

Qc ¼ 34f 0:52

which implies a smaller frequency dependence,0.52 against 0.70, and a slightly greater Qo value,34 as confront of 25.

5.3. Comparison of the Q values for S-waves

In the present study we have obtained severalQ values for S-waves, observing that Q is stronglydependent on the method. We can express theQ-factor of the medium as the sum of two e¡ects,the near-surface attenuation just below the seismicstation, let us name it QU , and the seismic attenu-ation of the remain path of the seismic waves, letus call it Qpath. In general we could express thecombined contribution of both as:

DWQ31T ¼ D1 Q31

UþD2 Q31

path ð10Þ

where D is the total path, D1 is the thickness ofthe shallowest layer where the near-surface at-tenuation occurs, and D2 is D3D1. We shouldexpect that the QU value is smaller, higher attenu-ation, than the Qpath value, and the combinedcontribution, QT, would be closer to the QU value.The di¡erent QL values indicate that the codanormalization method provides higher estimationsthan the spectral or the broadening-of-the-pulsetechniques. From this observation, we can assumethat the spectral or broadening-of-the-pulse meth-ods measure directly QT, but the coda normaliza-tion method or the single back-scattering modelprovide Qpath. This a⁄rmation is based on theprocedure followed by the methods. The ¢rstmethod performs the ratio between the attenua-tion of the direct S-waves and the coda waves.Both types of waves have been recorded in thesame station, and they have the same contributionof the near-surface attenuation, QU , and the ratiocancels this near-surface attenuation. Therefore,the measured Q value only re£ects the attenuationof the path, Qpath, and it is greater than QT . A

similar idea is applied for the coda-waves method.Using this method we estimated the decay of thecoda amplitude with time. Every point of thecoda amplitude has both Q contributions, butQU is the same for all the points of the codaenvelope. When we measure the decay ratio ofthe coda envelope, we eliminate the near-surfaceattenuation e¡ects and obtain only the measure ofthe path attenuation. As an exercise we can esti-mate QU for the array site, comparing the Q ob-tained by the coda normalization method,Qd = 140, and the Q from the spectral S-wavesmethod, QL = 41, under the assumption that QL

is the total Q, QT . We shall use as reference thefrequency of 15 Hz, where both values are valid.Thus,

Q31L

¼ Q31U

þQ31d ð11Þ

and substituting values we obtain that QU = 58,the value of the near-surface attenuation. ThisQ, for the interval distance used, 1^5 km, impliesa U of 0.023, a normal value for a zone with highattenuation, as any volcanic area.

5.4. Separation of intrinsic and scatteringattenuation

It is possible to study the contribution of theintrinsic and scattering attenuation over the mea-sured Q value of the seismic waves under the hy-pothesis of:

Q31T ¼ Q31

S þQ31i ð12Þ

where QS represents the scattering Q, and Qi isthe intrinsic Q. In the literature there are twomain techniques that permit this separation; the¢rst one was developed by Wennerberg (1993)and the second one is the Multiple Lapse TimeWindow Analysis (MLTWA), described by Hoshi-ba (1993). Both of them have been widely used inseveral regions and in the present work we shalluse only the ¢rst one, because the MLTWA meth-od cannot be applied in the present distancerange. The MLTWA method is based on the com-putation of the energy integrals on (generally)three successive time windows (Sato and Fehler,1998). Due to the low values of the S-wave veloc-ity (1.5 km/s), as well as the short duration of the



seismic signals that implies the use of very shorttime windows (3 s), the integral that de¢nes theenergy density in each time window cannot besolved on a su⁄cient number of points. Thisfact involves that the energies of each time win-dow cannot be separated one from each other forthe complete range of variability of both the seis-mic albedo and the total attenuation (B0 =Rs/(Ri+Rs) and Le31 =Ri+Rs, respectively). Conse-quently, for the seismic signals recorded at Decep-tion the MLTWA method cannot be applied.

Following the method described by Wenner-berg (1993), in the 3-D case, Qc

31 can be ex-pressed as a function of intrinsic and scatteringattenuation parameters by:

Q31c ¼ Q31

i þQ31S ð132 N ðQ ÞÞ ð13Þ

where the function N(Q) was derived by Wenner-berg (1993) using a numerical ¢x, given:

132N ðQ Þ ¼ 31=ð4:44 þ 0:738Q Þ ð14Þ

with:

Q ¼ Q31S g t ð15Þ

where g is the angular frequency and t is the lapsetime. The relationships of Eqs. 12 and 13 can beused as a system of equations to solve for Qi

31

and QS31.

To apply the above described method we needto estimate both QT and Qc. In our case we havefor S-waves three possible estimations of QT , i.e.the spectral Q, the broadening of the pulse Q andthe coda normalization Q. Spectral Q and thebroadening of the pulse Q are similar and we shalluse the spectral value. This is due to the uncer-tainty we have with the constant c used in theshear-waves broadening pulse method. Althoughwe have in the previous section noted that Qd

does not represent the total attenuation, we shalluse also this value, as a test, to separate the in-trinsic and scattering attenuation. The abovemethod requires an additional condition, theequation system can be solved if QT is smallerthan the Qc value, implying that for Qd the validinterval of frequency is the 13^25-Hz band.

In Fig. 17a we represent the results of this sep-aration using the Qd values. As observed, Qs isgreater than Qi, and Qi is close to the Qc estima-

Fig. 17. Qi and Qs values derived as a function of the Q of S-waves used. (a) Using the coda normalization method; (b) usingthe spectral Q.



tions. The fact to obtain this so high Qs value,and greater than Qi, is not in agreement withothers results obtained in di¡erent volcanic areas,as for example Etna and Campi Flegrei (Del Pez-zo et al., 1995) or Vesuvius (Bianco et al., 1999).It is di⁄cult to image that in a so heterogeneousarea, as this volcanic environment, the scattering-Q could have a value as high as reported in thisseparation. This unrealistic result indicates thatthe use of the present Qd value for the separationof Qi and Qs cannot be correct. It may be due tothe fact that our Qd does not take into accountthe near-surface attenuation, which for thepresent short distance range is very important.Therefore, it seems to be more correct to use theother Q values derived for S-waves. For example,using the spectral QL , we separate the intrinsicand scattering e¡ects in the 7^15-Hz band, wherethe spectral QL and Qc are valid simultaneously(Fig. 17b). We observe that intrinsic-Q is verysimilar to coda-Q, and scattering attenuation isalways stronger than intrinsic attenuation, withthe exception at 7 Hz. These results indicatethat, for Deception Island volcano, the scatteringe¡ect is predominant over the intrinsic attenua-tion, as could be expected in such a heterogeneousarea. This result is in agree with other areas asEtna volcano (Del Pezzo et al., 1995) or Vesuviusvolcano (Bianco et al., 1999). This result seems tobe the expected one due to the strongly heteroge-neous structure of Deception Island volcano, re-£ected in many e¡ects such as the distortion ofthe wave ¢eld observed by Saccorotti et al. (2001)or the complex fracture pattern described by Iba¤-n‹ez et al. (2003).

5.5. Shear-wave splitting observations

Beyond the e¡ect over the amplitude of theseismic waves of the heterogeneous structure ofDeception Island volcano, which provides a great-er scattering attenuation than the intrinsic e¡ects,we observe other phenomena, including a clearalignment of the polarization of the fast S-wavesthat may suggest the presence of seismic anisotro-py. Shear-wave splitting in volcanic areas hasbeen evidenced at Long Valley Caldera (Savageet al., 1990), at Hawaii (Booth et al., 1992;

Savage et al., 1989), at Vesuvius (Bianco et al.,1998a), at Mt. Etna (Bianco et al., 1998b), andat Mt. Ruapehu (Miller and Savage, 2001), gen-erally evidencing a correlation between seismicanisotropy and the local or regional maximumprincipal stress orientation acting on the analyzedareas. When a shear-wave enters an anisotropicvolume, it splits into two phases each propagatingwith di¡erent polarizations and velocities. Thesetwo phases (qS1 and qS2) arrive at the aniso-tropic/isotropic interface at di¡erent times, neverreconstructing the original waveform (Crampin,1981), de¢ning two measurable ‘splitting parame-ters’ that are the time lag between qS1 and qS2(hereafter TD) and the polarization direction ofthe qS1 phase (hereafter qS1PD). We checked forthe presence of shear-wave splitting for the seis-micity recorded at Deception during the January^February 1999 swarm. Among the 863 localizedevents we selected 80 volcano^tectonic earth-quakes recorded at the three 3-D stations accord-ing to the following criteria: (1) S-wave phaseswith high signal to noise ratios, and (2) incidenceangles strictly inside the theoretical estimated val-ue of shear-wave window, in order to avoid thefree-surface interaction that may corrupt theS-wave polarization (Booth and Crampin, 1985).The shear-wave window is de¢ned by a criticalangle ic = sin31vS/vP ; it is 37‡ in a half spacewith a Poisson’s ratio of 0.25.

We shall discuss only on the results we obtainedmeasuring the qS1PD, as our TD measurementsare not reliable, show a lot of scatter, and needmore appropriate investigations that are beyondthe scope of the present work. We measured theqS1 polarization direction using a quantitativemethod based on a 3 by 3 particle motion cova-riance matrix decomposition (Jurkevics, 1988).The qS1PD is the direction of the horizontal pro-jection of the eigenvector (related to the largesteigenvalue) obtained by diagonalizing the cova-riance matrix. This method has the advantage ofcomputing the eigenvector using the 3-D wave-form even if the representation is 2-D, constrain-ing more stable results in very heterogeneous me-dia. The qS1 polarization directions show scatter,but a prevalent E^W striking direction is evidentat all the stations (Fig. 18).



The pattern of the observed splitting parameteris compatible with the presence of an anisotropicvolume probably con¢ned to the last km of theupper crust. This anisotropic volume is mainlysuggested by the preferred orientation of qS1PD

that has been correctly estimated and is also themore stable parameter. However, due to the lackof information on both structural features andstress ¢eld geometry, detailed discussion of thefeatures of the suggested anisotropic volume isbeyond the scope of the present work. We onlywant to emphasize the presence of anisotropy tomake a correlation with the observed scatteringproperties of the upper crust.

6. Conclusions

Using di¡erent techniques and P-waves, S-wavesas well as coda-waves, we estimated the seismicattenuation of Deception Island volcano in the6^30-Hz frequency interval, and for short hypo-central distances. In general, results reveal thatthis environment has a high attenuation; a con-clusion derived from the observed low Q values,as in comparison to other volcanic or tectonicareas. Methods used in the present work provideus frequency-independent and -dependent Q val-ues, as well as the analysis of a single pulse orpackage of waves. We observed a high ratio QP/QL , between 1.9 and 3, and also an anomalousfrequency dependence of Q for S- and coda-waves. This type of results is not di⁄cult to ¢ndin other volcanic environments, where the com-plex surface structure a¡ects strongly the propa-gation of seismic waves. Another important ob-servation is the di¡erence found of Q for S-wavesas a function of the method used. We interpretedthe di¡erences among Q for the same type ofwaves as an e¡ect of the elimination of thenear-surface attenuation e¡ect due to the methodof calculus. The coda normalization and the singleback-scattering methods do not provide informa-tion about the near-surface attenuation. However,the spectral or the broadening-of-the-¢rst-pulsemethods introduce this e¡ect in the Q values.Under this assumption, we separated Q, forS-waves, into two contributions, the near-surfaceattenuation, and the remaining Q. For 15 Hz, thenear-surface Q is 58, and 140 for the other por-tion of the structure not deeper than 5 km, whichreveals a high contrast. To investigate the e¡ect ofthe heterogeneous structure of Deception Island

Fig. 18. The predominant eigen-directions derived from thepolarization analysis, for the three 3-D stations of the seismicarray.



over the seismic attenuation, we separated the in-trinsic and scattering Q for S-waves. Results in-dicate that the scattering process is predominantover the intrinsic attenuation, and that the Qc andQi values are very similar, at least in the fre-quency interval used. This scattering e¡ect, pro-duced probably by the complex structure of theisland, also is re£ected in the splitting phenomenaof the S-waves, with evident delays in the arrivalof the horizontal components, longer than 0.1 s,for such a short hypocentral interval.

The evidences of the complex surface structureof Deception Island volcano have been observedalso in other studies, as for example Iba¤n‹ez et al.(2000, 2003) or Saccorotti et al. (2001). More de-tailed analyses are necessary to de¢ne the realstructure of the island based principally on tomo-graphic techniques, velocity and attenuation.Also, a deeper study about the real e¡ect of thestructure over the splitting of the shear-waves isnecessary.

Acknowledgements

This work was partially supported by projectsANT98-1111, REN-2000-2897 and REN-2001-3833. It was signi¢cantly improved by MarthaSavage and an anonymous reviewer. We thankE. Carmona, J. Morales, G. Alguacil, and M.Abril for useful comments on the manuscript.Thanks are due also to the sta¡ and scientists ofthe Gabriel de Castilla Base. Logistic support wasprovided by the Spanish ‘Eje¤rcito de Tierra’ and‘Armada’.

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