Lost tsunami

Maria Teresa Pareschi,1 Enzo Boschi,1 and Massimiliano Favalli1

Received 4 August 2006; revised 6 October 2006; accepted 18 October 2006; published 28 November 2006.

[1] Numerical simulations support the occurrence of acatastrophic tsunami impacting all of the easternMediterranean in early Holocene. The tsunami wastriggered by a debris avalanche from Mt. Etna (Sicily,Italy) which entered the Ionian Sea in the order of minutes.Simulations show that the resulting tsunami waves wereable to destabilize soft marine sediments across the IonianSea floor. This generated the well-known, sporadicallylocated, ‘‘homogenite’’ deposits of the Ionian Sea, and thewidespread megaturbidite deposits of the Ionian and SirteAbyssal Plains. It is possible that, �8 ka B.P., the Neolithicvillage of Atlit-Yam (Israel) was abandoned because ofimpact by the same Etna tsunami. Two other Pleistocenicmegaturbidite deposits of the Ionian Sea can be explainedby previous sector collapses from the Etna area.Citation: Pareschi, M. T., E. Boschi, and M. Favalli (2006),

Lost tsunami, Geophys. Res. Lett., 33, L22608, doi:10.1029/

2006GL027790.

1. Introduction

[2] In the early Holocene the eastern flanks of Mt. Etnavolcano (Figure 1a) suffered a large sector collapse(s), toform the Valle del Bove scar [Calvari and Groppelli, 1996;Calvari et al., 1998, 2004]. At the mouth of the Valle delBove, the Milo debris avalanche deposit [Calvari et al.,1998] stands above the Chiancone fan, the Chiancone beingevident from a seaward bulge in the coastline [Calvari andGroppelli, 1996; Del Negro and Napoli, 2002]. Offshore ofChiancone, a debris avalanche deposit, 250 km2 in area and20 km long, mantles the irregular sub-marine slopes[Pareschi et al., 2006a]. The total volume of the land-based(below sea level) Chiancone [Calvari and Groppelli, 1996]and offshore deposits [Pareschi et al., 2006a] is about 20–25 km3. Calvari et al. [2004] propose that the Chianconedeposits were emplaced, mostly, during a single collapseepisode. Likewise the offshore debris avalanche depositswere probably emplaced in a single collapse [Pareschi etal., 2006a]. If sufficiently rapid, such a catastrophic land-slide entering the sea may be expected to generate asignificant tsunami impacting low lying coastal regions ofthe eastern Mediterranean. However, a 10-m increase in sealevel over the last 8 kyr [Nir, 1997] has seriously compro-mised our ability to identify these in situ tsunami deposits.In addition, the inland and offshore deposits related to alandslide-triggered tsunami often comprise reworked mate-rials of different ages, which complicate attempts to make astraightforward event reconstruction.[3] Here we have used numerical simulations to constrain

the problem. First of all, we have simulated a reference

tsunami: Tsunami-I. This model tsunami is triggered by adebris avalanche from Mt. Etna, with a volume of 25 km3

and an initial velocity of 100 m/s. We next assess the impactof Tsunami-I across the Mediterranean Sea, before deter-mining the capabilities of Tsunami-I to destabilize softmarine sediments and to trigger sediment flows filling theIonian Basins. We finally look for field evidence of theseinstabilities and discuss them within the framework ofTsunami-I.[4] Such an event occurring today would impact signif-

icant coastal zones; thus from a hazard point of view it isessential to understand the potential impact of such acatastrophic landslide entering the sea. Our aim has thusbeen to perform an analysis of the likely effects of a largeMt. Etna debris avalanche and its ability to trigger cata-strophic tsunamis in the eastern Mediterranean.

2. Tsunami Simulation

[5] We simulate the tsunami propagation using a numer-ical model that takes into account nonlinearity, dispersion,refraction, diffraction, wave-to-wave interaction, bottomfriction and run-up [Wei et al., 1995; Kirby et al., 1998;Chen et al., 2000]. We also utilize a source model thatsimulates the effects of a landslide entering the sea to set upthe initial conditions for water displacement and tsunamivelocity [Pareschi et al., 2006b]. We use a Lambert Con-formal Conic Projection for the simulation domain (Spheroidand Datum: WGS84; 1st and 2nd Standard Parallel: 43�Nand 32�N) and take our bathymetry from the GEBCOdatabase (available at http://www.ngdc.noaa.gov/mgg/gebco/grid/1mingrid.html), providing bathymetry at 1 arcminuteresolution. In the Ionian Sea, a re-sampled computationalgrid with a 500 m step is used.[6] In our simulations, landslide volumes range from 5 to

30 km3, the upper bound being approximately the volumeof the on-shore (below sea level) Chiancone deposit plus theoff-shore debris avalanche deposit. Numerical initial land-slide velocities range from 100 to 25 m/s (average velocities50 m/s and 12.5 m/s respectively), implying travel times of400 to 1600 s for a uniformly decelerating landslideextending 20 km offshore (i.e. this is the time required tocover the maximum extent of the offshore debris avalanchedeposit, given our modelled velocities).[7] Our case-type tsunami (Tsunami-I) considers a 25 km3

landslide entering the sea with an initial velocity of 100 m/s.Tsunami-I affects the whole eastern Mediterranean Basin,including the coasts of Southern Europe, North Africa andAsia Minor (Figure 1 and Animation S1 of the auxiliarymaterial1). A train of waves, induced by phase dispersion[Ward, 2001], develops in the Central Ionian Sea and

1Auxiliary materials are available in the HTML. doi:10.1029/2006GL027790.

GEOPHYSICAL RESEARCH LETTERS, VOL. 33, L22608, doi:10.1029/2006GL027790, 2006ClickHere

for

FullArticle

1Istituto Nazionale di Geofisica e Vulcanologia, Pisa, Italy.

Copyright 2006 by the American Geophysical Union.0094-8276/06/2006GL027790$05.00

L22608 1 of 6

Central Levantine Basin (Figure 1, see also Animation S1and Text S1 of the auxiliary material). Characteristicwavelengths (l) are some tens of kilometres. Wavelengthsof around 50 km in the Central Ionian Sea (water depthh � 3,500 m, position CR in Figure 1) decrease to about20 km in the Southern Aegean at h � 500 m. The periodis of the order of minutes.[8] We also observe a number of effects induced by local

and regional topography. Between Africa and Crete, thewaves are observed as crushing together, a result of part ofthe wavefront rebounding off of the African coast. Inaddition, submerged flat topped highs (south of Zakynthosand south-west of Crete) enter into resonance with thetsunami period (Figure 1, Animation S1). The parabolicshape of the Sirte Gulf also induces wave focusing a few tensof kilometres offshore of the North African coast (Figure 1,Animation S1), and shallow continental shelves and sea-floorhighs induce wave height increases. In the northern Aegean,tsunami penetration is inhibited by multiple, scattered islands(average island-to-island distance = 30 km, comparable withthe tsunami wavelengths in that sea, see Figure 1).[9] Computed wave run-up heights reach maximum

values of about 40 m along the Calabrian coasts (SouthItaly), just to the North the landslide area (Figure 1).Because these coasts are steep and rocky, they are notsuitable for preserving tsunami deposits. In the simulation,those coasts of Greece and Libya (Africa) that face thelandslide area are impacted by wave run-ups 8–13 m.Computed run-ups reduce to 2–4 m on the coasts ofEgypt and Israel-Syria (including the area of Atlit-Yam,see Figure 1). Near-coast topography/bathymetry couldhave been locally largely changed during the last 8 ka.In addition the grid step, computationally adequate to

resolve the waveform in open sea, is very rough for run-up evaluation. As a consequence, the numerical run-upswe have here computed have to be considered veryapproximate estimations [Pareschi et al., 2006b].

3. Sea-Floor Instabilities Triggered by TsunamiPassage

[10] In the Ionian Sea, Tsunami-I impact is predicted tobe quite severe. According to Pareschi et al. [2006b], weexpect that tsunami-induced pore overpressure liquefies softsea-bed sediments. Soft marine sediments are stable (do notliquefy) when exposed to a tsunami wave of crest heightDa, if:

Da < C= rw g sin qð Þ ð1Þ

where q is the sea-bed slope angle, C is cohesion, rw iswater density and g is acceleration due to gravity [Pareschiet al., 2006b]. This relationship holds for linear long-waveapproximation [Ward, 2001] by applying the Mohr-Coulomb criterion [Hampton et al., 1996]. It presupposesan incompressible (negligible gas content) and elasticmedium (without hysteresis effects). In any case, in oursimulations the first wave displays the maximum crestheight of all the waves in the train. Moreover, duringincipient failure, cohesive bonds are gradually broken.Relation (1) thus provides a first order estimate for stableconditions of marine soft sediments when exposed to thepassage of a tsunami.[11] If we consider the upper 10 m of sediments in cores

taken from sites across eastern Mediterranean (Ocean Dril-ling Project (ODP) leg 160), which mainly consist ofnannofossil and silt ooze plus Sapropels [Kopf et al.,

Figure 1. (a) Maximum wave crests heights predicted by the Tsunami-I scenario in the eastern Mediterranean. Blue linesare arrival times, s, of the first tsunami maximum. CR, MR, CMR, P1, P2, and P3 are some locations discussed in the text.Tsunami-I wave field at time steps of (b) 1 h 200, (c) 2 h, and (d) 3 h 200 after the entrance of the debris avalanche into thesea. The white vertical bars provide the tsunami wave height scale.

L22608 PARESCHI ET AL.: LOST TSUNAMI L22608

2 of 6

1998], we obtain yield strengths versus depth relationshipswhich give typical values for cohesion (= yield strength atthe sea floor) of the order of kPa. For example, C of around2.5 kPa is obtained at site ODP 160–964-A, at a depth of3650 m. Similar values are also found by Sultan et al.[2004] in undrained conditions, at continental marginswhere a 10 – 20% sand fraction is present.[12] Following relationship (1), Figure 2 shows those

areas where instabilities will likely to be triggered by thepassage of Tsunami-I, given soft marine sediments mantlingthe sea floor and a value of 1 kPa for C.

4. Evidence for Instabilities From a PastEtna-Collapse-Derived Tsunami

[13] Field-evidence for sediment re-working and/or failureare apparent in the Ionian Sea. These occur at five main

locations (Figure 2). The first comprises the widespreadhomogenite deposits that characterise the cobblestone topo-graphy of the Calabrian Ridge [Kastens and Cita, 1981] as,for example, at location CR (Figures 1 and 2). The secondlocation includes the homogenite deposits of the Mediter-ranean Ridge cobblestone topography [Cita et al., 1996, andreferences therein], as, for example, location MR (Figures 1and 2). The third is the Augias megaturbidite deposits of theIonian and Sirte Abyssal Plains [Hieke and Werner, 2000,and references therein] (IAPM and SAPM, Figure 2). Thefourth and fifth are deposits in the Western HerodotusTrough [Rebesco et al., 2000, and references therein](WHT, Figure 2) and along a transect in the MatapanTrench [Rebesco et al., 2000] (MT, Figure 2), respectively.[14] Because these deposits are spread across a wide

region and are topographically disjointed, they have beencorrelated to a tsunami triggered by the Late Bronze Ageeruption of Santorini [Kastens and Cita, 1981; Cita et al.,1996; Hieke and Werner, 2000]. Recently, however, on thebasis of numerical simulations, Pareschi et al. [2006b] haveruled out any relationship between these marine depositsand the Santorini eruption. Instead, however, we can relatethem to instabilities induced by a tsunami source at Mt.Etna. The homogenite of the closed, pond-like basins of thecobblestone topography of the Mediterranean and CalabrianRidges is located at the basin floors, with maximum thick-nesses of several metres. It originated from the steep slopesof the basins themselves. At site CR (Figures 1 and 2), byinterpolating Kastens and Cita’s [1981] morphological data,we find that 1–7% of the basin’s slopes exceed 27� (thisbeing the internal friction angle for nannofossil/silt ooze[Kopf et al., 1998]). Forty-to-fifty percent of the slopes alsoexceed 11�, which is the internal friction angle of theSapropel S1 layer [Kopf et al., 1998]. The name ‘‘homog-enite’’ was assigned by Kastens and Cita [1981], because,except for their basal part, such marine deposits are quitehomogeneous, with fine silt to clay grain-sizes and nograding. The clay content dominates over silt, in contrastto normal soft marine sediments (ODP Leg 160–964 [Kopfet al., 1998]). These homogeneous characteristics are relatedto tsunami dynamics (see Text S2 of the auxiliary materialfor further details). Homogenite overlays a highly variable,few decimetres-thick, normal deposition layer and then theSapropel layer S1-A [Kastens and Cita, 1981]. In theregion, the marker S1 is 14Cnc dated at its top at about8 ka B.P. [Troelstra et al., 1991; Hieke and Werner, 2000,and references therein].[15] Tsunami-I would also have been capable of trigger-

ing bed sediment instabilities in the basins of the Calabrianand Mediterranean Ridges, thus explaining the occurrenceof homogenite in these locations. In fact, at site CR forexample, the transitional area between normal (hillslope)pelagic areas and eroded homogentite-source slopes havemaximum values of 2–3� (from Kastens and Cita’s [1981]data). Given these values and, as already stated, by hypoth-esizing C = 1 kPa in relationship (1), tsunami crest heightsof greater than 2–3 m will induce instability (see alsoFigure 3a). As a confirmation, at site CR, the eroded areashave an average slope of 10� and Tsunami-I has a maximumwave crest height of 6 m (Figure 3a). Such an event wouldthus have been capable of inducing slope instabilities to

Figure 2. Map of the Ionian Sea showing those areasacross which Tsunami-I potentially induces instabilities (redareas). Bathimetric cyne-line interval is 1000 m. Yellow linelabelled ECA identifies the External Calabrian Arc. The pinkline (IGW) locates the main high (watershed) of the SirteRise, terminating in the Ionian Gap. P1, P2, and P3 markthree unstable areas within the Sirte Gulf. Orange linesmark some maximum slope drainage paths. Two yellow dotsmark turbidite deposits in the Calabrian (CR) and Medi-terranean (MR) Ridges [after Cita et al., 1996]. Note that atCMR no homogenite was found [Kastens and Cita, 1981].MER and MESC16 locate transects by Marani et al. [1993]and Argnani and Bonazzi [2005]. IAPM and SAPM whitelines equal Augias deposits [after Hieke and Werner, 2000];LC14 is the core in the WHT deposit (Rothwell et al.,unpublished manuscript, 1995). MT is the transect within theturbidite deposit of the Matapan Trench [after Rebesco et al.,2000]. Finally, the red dot locates ODP leg 160-964 and SN-1is the marine station.

L22608 PARESCHI ET AL.: LOST TSUNAMI L22608

3 of 6

form the observed eroded areas and homogenite deposits ofthe Calabrian Ridge.[16] At site MR on the Mediterranean Ridge (Figure 2),

Tsunami-I had a modelled crest height >2 m, and wasthereby capable of inducing instabilities on slopes �2–3�.This is consistent with widespread occurrences of homog-enite deposits on the basin floors across this area [Cita et al.,1996].[17] In contrast, at site CMR (Figure 1), no traces of

homogenite were found in situ [Kastens and Cita, 1981].CMR is located on a plateau with average slopes <0.05�.Relationship (1) is consistent with no bed sediment desta-bilization at this site, during the passage of Tsunami-I(tsunami wave heights �1.8 m, Figure 3c).[18] The Augias megaturbidite [Hieke and Werner, 2000]

covers the Ionian and Sirte Abyssal Plains (IAPM andSAPM in Figure 2). IAPM and SAPM are separated bythe Ionian Gap (end-point of the IGW watershed in Figure 2).Another ‘‘turbidite’’ deposit occurs in the Western Herodo-tus Trough (WHT). The shallow water provenience ofIAPM and SAPM was inferred by the significant presenceof fragments of the shallow-water green alga Halimeda[Hieke and Werner, 2000]. By including the MT deposittoo, the total volume of all the deposits is 165 ± 10 km3

[Rebesco et al., 2000].[19] Characteristics of IAPM are an upper body that is

practically homogeneous in terms of grain-size (silt-clay)and composition, and basal coarser turbiditic beds whichoccur above the Sapropel S1 layer, this last unit beinglocally (radiocarbon) aged as 9.1 ± 0.7 ka [Hieke andWerner, 2000]. Also SAPM and WHT are characterizedby an upper homogeneous layer and by a lower coarser one.Dating of those deposits is however not available (R. G.Rothwell et al., Deacon Laboratory, unpublished manu-

script, 1995, hereinafter referred to as Rothwell et al.,unpublished manuscript, 1995).[20] Our Tsunami-I modelling indicates that the African

slopes of the Sirte Gulf are likely sources for IAPM andSAPM. In particular, our modelling shows that Tsunami-Iwould have been able to generate some unstable zones inthis area due to the shallow waters at the transitionbetween continental slopes and shelves, wave focusing,local sea-floor highs, and the sloping seafloor. As anexample Figures 3d–3f show Tsunami-I maximum wavecrest heights at locations P1, P2 and P3 in the Sirte Gulf(Figure 2). At these localities, following relation (1), thewave crest heights (2–5 m) of Tsunami-I are able toinduce sediment liquefaction.[21] The destabilized sediments are potentially able to

flow downslope into SAPM and IAPM. In fact, the GEBCObathymetry shows smooth, regular maximum-slope drainagepaths, hundreds of kilometres long, developing from thedestabilized areas near the African coasts and ending atIAPM and SAPM (Figure 2). In particular, while the westernportion of the Sirte Gulf feeds IAPM, the eastern Sirte slopesfeed SAPM (Figure 2). However, because the maximumslope paths bend at the Sirte Rise, where the flows display anerosive flow passage [Hieke and Werner, 2000], sedimentflows could have overturn the IGW watershed by inertia.This suggests volume budgets between SAPM and IAPMcannot be simply performed on the basis of destabilizedareas of the Sirte western/eastern slopes. The long pathexperienced by these marine sediments explains their finegrain size. Such long sediment flow paths are not unusual.For example, the 1929 Grand Banks earthquake triggered asediment flow which flowed on gentle (1–0.03�) slopes, withvelocities of 30 to 5 m/s, for well over 800 km, mobilizing atotal volume of 200 km3 [Heezen and Ewing, 1952].

Figure 3. Simulated maximum tsunami crest heights (red lines + numbers in m) as function of the volume and initialvelocity of a debris avalanche entering the sea from Mt. Etna, at sites (a) CR, (b) MR, (c) CMR, (d) P1, (e) P2, and (f) P3(see Figure 2 for locations). Homogenite was found at sites CR and MR, but not at CMR [Kastens and Cita, 1981]. Bluecircle marks Tsunami-I. Sites P1, P2, and P3 are predicted to be unstable areas of the Sirte Gulf when impacted by Tsunami-I.At those sites (Figures 3d–3f), blue dotted lines mark the transition between tsunami heights capable of non-inducing andinducing sediment instabilities. Any landslide volume-velocity pair above the line is capable of inducing sediment instabilities.

L22608 PARESCHI ET AL.: LOST TSUNAMI L22608

4 of 6

[22] According to relationship (1), Tsunami-I suggeststhat contributions to WHT are expected from the Africanslopes and to MT from the nearby instable slopes (Figure 2).[23] We have also looked for offshore evidence (in the

literature) of a tsunami passage in the Ionian Sea, betweenCalabria and Sicily, beyond the area covered by the debrisavalanche deposits [Pareschi et al., 2006a] (Figure 2). Here,it is very difficult to seismically identify traces of a tsunamipassage, because the related deposits are expected to havecoarser compositions with respect to the homogenite andupper Augias deposits (associated with easily identifiableseismically transparent layers [Rebesco et al., 2000]). Inaddition, near Sicily and Calabria, the finest fraction of thedebris avalanche deposits probably contributed to the for-mation of turbidity currents, and relationship (1) becomesinadequate to check the conditions for the onset of sea-bedsediment instabilities.[24] However, there is potential tsunami evidence off-

shore the Southern Calabria and eastern Sicily coasts. Manygravity sliding fronts, tens of kilometres long, occur off-shore from the Southern Calabria coasts [Rossi and Sartori,1981] (north-west portion of the ECA area in Figure 2).Because of the low resolution of the survey, however it isimpossible to discriminate potential tsunami traces fromtectonic gravity processes.[25] Offshore of eastern Sicily, at the foot of the Malta

Escarpment, at a depth of 2300–2500 m, average slopes are1�, and have a N-S orientation (Figure 2). At this location,asymmetric, climbing bedforms occur [Marani et al., 1993](transect MER of Figure 2). The sequence, with wave-lengths of a few kilometres, amplitudes of 25 to 35 m and atotal thickness of 250 m, is discontinuous. It has beenrecognized as being related to turbidity activity, as well asto water currents [Marani et al., 1993]. Argnani andBonazzi [2005] identified eroded sedimentary strata, at awater depth of 3000 m, south of MER (profile MESC16 ofFigure 2). Because no strong bottom currents have beenrecorded offshore from eastern Sicily (INGV-INFN marinestation SN-1 of Figure 2, available at http://www.garr.it/conf_06/presentazioni/GARR_CataniaNEMO-SN1.pdf),bothMarani et al. [1993] and Argnani and Bonazzi’s [2005]data could possibly be interpreted as traces of a tsunamiimpact.[26] Figures 3a–3f show maximum wave crest heights at

some sites in the Ionian Sea, as a function of the volume andthe initial velocity of a Mt. Etna landslide entering the sea.By taking into account both relation (1) and local slopes atthe sites of Figures 3a–3f, and the offshore-detected Mt.Etna-landslide volume of 20–25 km3 [Pareschi et al.,2006a], we infer that the velocity of the Early HoloceneMt. Etna debris avalanche was a few ten meters per second(minimum). Triggered by this landslide, Tsunami-I is able toexplain homogenite, Augias, WHT, and MT deposits of theIonian Sea.[27] According to Figures 3a–3f, many possible vol-

umes and velocities of a Mt. Etna landslide can triggerinstabilities that result in such deposits. However, relation-ship (1) and local slopes suggest velocities must be greaterthan 25 m/s and volumes above 10 km3. A tsunami(smaller than Tsunami-I), triggered by a 10 km3 landslideentering the sea at 50 m/s would induce 25 m run ups

along the Calabrian coasts and 1–2 m run-ups along theIsraeli coasts.

5. Age Correlation

[28] The upper portion (up to 30 m thick) of the Chian-cone deposits formed after 7.600 ± 130 (rad. cal.) yr B.P.[Calvari and Groppelli, 1996]. A soil fragment foundupslope of the Chiancone [Calvari et al., 1998], is dated>8.4 ka B.P.; but it could be a reworked material (Calvari,private communication, 2006). Homogenite and Augiasdeposits overlay the Sapropel layer S1-A. We infer thatthe occurrence of Mt. Etna tsunami occurred between theend of S1-A and 7.6 ka B.P. The tsunami probably occurredsome time after the Sapropel S1 emplacement, because,between homogenite and Sapropel S1-A, there is a normalpelagic depositional layer [Kastens and Cita, 1981] (seealso Text S3 of the auxiliary material).[29] The submerged (10 m b.s.l.) village of Atlit-Yam is a

well preserved Pre-Pottery Neolithic settlement somehundreds meters offshore from the current coasts of Israel(Figure 1). This site shows evidence of sudden abandon-ment about 7.6 ka B.P. (not.cal). For example, below a claystratum, a pile of fish gutted and processed in a size-dependent manner and then stored for future consumptionand trade has been discovered [Zohar et al., 2001]. TheHolocene sea level rise in Mediterranean [Nir, 1997] cannotsimply explain this. We thus ask if Atlit-Yam was aban-doned as the result of the Etna collapse which then triggereda tsunami with our Tsunami-I characteristics. Future analy-ses are desirable to search for tsunami depositional featuresat Atlit-Yam.

6. Conclusions

[30] In the early Holocene, a flank failure of Mt. Etnagenerated an on-offshore debris avalanche deposit trigger-ing a large tsunami [Pareschi et al., 2006a, and referencestherein]. On-shore tsunami deposit evidence from this eventis lacking. Instead, we find (from literature) a wealth of sub-marine evidence. Numerical simulations indicate that thecollapse occurred as a rapid and catastrophic event so thatthe resulting tsunami impacted all the eastern Mediterra-nean. The Holocene marine deposits of the Ionian Sea canbe explained by such an event. Other causes for thetsunami can be discounted. It could not, for example,have been an asteroid-tsunami because two megaturbiditelayers underlay the Augias deposit (see next paragraph). Itcould also not have been an earthquake-tsunami becausehigh wave amplitudes must occur in the deep CentralIonian Sea to trigger homogenite. As a confirmation aboutthis point, in the last 2.0 ka, some very strong earthquakesand related tsunamis have affected the Ionian Sea (Text S4of the auxiliary material), without triggering there anyfloor destabilization.[31] In the Ionian Sea, beneath the Augias deposits and

inter-bedded by seismically reflective strata, are two othersimilar megaturbidite layers, their volumes are respectivelysmaller and greater than the Augias deposit [Hieke, 2000].At transect MT (Figure 2), below the upper-most deposit,another seismically transparent layer occurs; but data reso-lution does not allow detection of an additional deeper

L22608 PARESCHI ET AL.: LOST TSUNAMI L22608

5 of 6

transparent deposit [Rebesco et al., 2000]. We suggest thatthese further, older layers are suggestive of tsunami pas-sages across the eastern Mediterranean because they aresimilar to the upper transparent layer we have linked to theValle del Bove collapse, and the triggering tsunamis orig-inated by collapses in the Etnean area. Indeed, offshore ofMt. Etna, seismic data show evidences for distinct land-slides deposits, including a catastrophic, multiple, synchro-nous set of slumps [Pareschi et al., 2006a].[32] In the future, we will address whether secondary

tsunamis could also have been triggered by those marinesediment flows caused by the passage of the initial tsunamifrom the Mt. Etna collapse.

[33] Acknowledgments. We thank F. Mazzarini and A. Camerlenghifor fruitful discussions, and J. T. Kirby for the code FUNWAVE in the web.We acknowledge the thoughtful revision by A. J. L. Harris and ananonymous reviewer, greatly improving our manuscript. We also thankKristina Sine, GRL senior editor’s assistant, for her thoughtful assistanceand professionalism. Funding was provided by the Italian Dipartimento diProtezione Civile to the Istituto Nazionale di Geofisica e Vulcanologia.

ReferencesArgnani, A., and C. Bonazzi (2005), Malta Escarpment fault zone offshoreeastern Sicily: Pliocene-Quaternary tectonic evolution based on new mul-tichannel seismic data, Tectonics , 24 , TC4009, doi:10.1029/2004TC001656.

Calvari, S., and G. Groppelli (1996), Relevance of the Chiancone volcani-clastic deposits in the recent history of Etna volcano (Italy), J. Volcanol.Geotherm. Res., 72, 239–258.

Calvari, S., L. H. Tanner, and G. Groppelli (1998), Debris-avalanchedeposits of the Milo Lahar sequence and the opening of the Valle delBove on Etna volcano (Italy), J. Volcanol. Geotherm. Res., 87, 193–209.

Calvari, S., L. H. Tanner, G. Groppelli, and G. Norini (2004), Valle delBove, eastern flank of Etna volcano: A comprehensive model for theopening of the depression and implications for the future hazards, inMt. Etna: Volcano Laboratory, Geophys. Monogr. Ser., vol. 143, pp.65–75, edited by A. Bonaccorso et al., AGU, Washington, D. C.

Chen, Q., J. T. Kirby, R. A. Dalrymple, A. B. Kennedy, and A. Chawla(2000), Boussinesq modeling of wave transformation, Breaking and run-up. II: 2D, J. Waterw. Port Coastal Ocean Eng., 126, 48–56.

Cita, M. B., A. Camerlenghi, and B. Rimoldi (1996), Deep-sea tsunamideposits in the eastern Mediterranean: New evidence and depositionalmodels, Sediment. Geol., 104, 155–173.

Del Negro, C., and R. Napoli (2002), Ground and marine magnetic surveysof the lower eastern flank of Etna volcano (Italy), J. Volcanol. Geotherm.Res., 114, 357–372.

Hampton, M. A., H. J. Lee, and J. Locat (1996), Submarine landslides, Rev.Geophys., 34, 33–59.

Heezen, B. C., and M. Ewing (1952), Turbidity currents and submarineslumps, and the 1929 Grand Banks earthquake, Am. J. Sci., 250, 849–873.

Hieke, W. (2000), Trasparent layers in seismic reflection records from thecentral Ionian Sea (Mediterranean)–Evidence from repeated catastrophicturbidite sedimentation during Quaternary, Sediment. Geol., 135, 89–98.

Hieke, W., and F. Werner (2000), The Augias megaturbidite in the centralIonian Sea (central Mediterranean) and its relation to the Holocene San-torini event, Sediment. Geol., 135, 205–218.

Kastens, K. A., and M. B. Cita (1981), Tsunami induced sediment transportin the abyssal Mediterranean Sea, Geol. Soc. Am. Bull., 89, 591–604.

Kirby, J. T., G. Wei, Q. Chen, A. B. Kennedy, and R. A. Dalrymple (1998),FUNWAVE 1.0: Fully nonlinear Boussinesq wave model documentationand user’s manual, Res. Rep. CACR-98–06, Cent. for Appl. Coastal Res.,Univ. of Delaware, Newark, Del.

Kopf, A., M. B. Clennel, and R. Flecker (1998), Relationship between thevariation of undrained shear strength, organic carbon content, and theorigin and frequency of enigmatic normal faults in fine-grained sedimentsfrom advanced poston cores from the Eastern Mediterranean, Proc.Ocean Drill. Program Initial Rep., 160, 645–661.

Marani, M., A. Argnani, M. Roveri, and F. Trincardi (1993), Sedimentdrifts and erosional surfaces in the central Mediterranean: Seismic evi-dence of bottom-current activity, Sediment. Geol., 82, 207–220.

Nir, Y. (1997), Middle and late Holocene sea-level along the Israel Medi-terranean coast–Evidence from ancient water wells, J. Quat. Sci., 12,143–151.

Pareschi, M. T., E. Boschi, F. Mazzarini, and M. Favalli (2006a), Largesubmarine landslides offshore Mt. Etna, Geophys. Res. Lett., 33, L13302,doi:10.1029/2006GL026064.

Pareschi, M. T., M. Favalli, and E. Boschi (2006b), Impact of the Minoantsunami of Santorini: Simulated scenarios in the eastern Mediterranean,Geophys. Res. Lett., 33, L18607, doi:10.1029/2006GL027205.

Rebesco, M., B. Della Vedova, L. Cernobori, and A. Aloisi (2000), Acous-tic facies of Holocene megaturbidites in the eastern Mediterranean, Sedi-ment. Geol., 135, 65–74.

Rossi, S., and R. Sartori (1981), A seismic reflection study of the externalCalabrian Arc in the northern Italian Sea, Mar. Geophys. Res., 4, 403–426.

Sultan, N., P. Cochonat, F. Cayocca, J. F. Bourillet, and J. L. Colliat (2004),Analysis of submarine slumping in the Gabon continental slope, AAPGBull., 88, 781–799.

Troelstra, S. R., G. M. Ganssen, K. van der Borg, and A. F. M. de Jong(1991), A late quaternary stratigraphic framework for eastern Mediterra-nean Sapropel S1 based on AMS 14C dates and stable oxygen isothopes,Radicarbon, 33, 15–21.

Wei, G., J. T. Kirby, S. T. Grilli, and R. Subramanya (1995), A fully non-linear Boussinesq model for surface waves. I. Highly nonlinear, unsteadywaves, J. Fluid Mech., 294, 71–92.

Ward, S. N. (2001), Landslide tsunami, J. Geophys. Res., 106, 11,201–11,215.

Zohar, I., T. Dayan, E. Galili, and E. Spanier (2001), Fish processing duringthe early Holocene: A taphonomic case study from coastal Israel,J. Archaeol. Sci., 28, 1041–1053.

�����������������������E. Boschi, M. Favalli, and M. T. Pareschi, Istituto Nazionale di Geofisica

e Vulcanologia, Via della Faggiola 32, I-56100 Pisa, Italy. (pareschi@pi.ingv.it)

L22608 PARESCHI ET AL.: LOST TSUNAMI L22608

6 of 6