the campi flegrei caldera: unrest mechanisms and hazards · rate of about 1.1–1.7 cm a−1,...

From: TROISE, C., DE NATALE, G. & KILBURN, C. R. J. (eds) 2006. Mechanisms of Activity and Unrest at LargeCalderas. Geological Society, London, Special Publications, 269, 25–45.0305-8719/06/$15.00 © The Geological Society of London.

The Campi Flegrei caldera: unrest mechanisms and hazards

G. DE NATALE1, C. TROISE1, F. PINGUE1, G. MASTROLORENZO1,L. PAPPALARDO1, M. BATTAGLIA2 & E. BOSCHI1.

1Istituto Nazionale di Geofisica e Vulcanologia, Via Diaclezano 328, 80124 Napoli, Italy(e-mail: [email protected])

2Department of Structural Geology & Geodynamics, University of Göttingen,37077 Göttingen, Germany

Abstract: In the last four decades, Campi Flegrei caldera has been the world’s most activecaldera characterized by intense unrest episodes involving huge ground deformation andseismicity, but, at the time of writing, has not culminated in an eruption. We present a carefulreview, with new analyses and interpretation, of all the data and recent research results. Wedeal with three main problems: the tentative reconstruction of the substructure; the modellingof unrest episodes to shed light on possible pre-eruptive scenarios; and the probabilisticestimation of the hazards from explosive pyroclastic products. The results show, for the firsttime at a volcano, that a very peculiar mechanism is generating episodes of unrest, involvingmainly activation of the geothermal system from deeper magma reservoirs. The character andevolution of unrest episodes is strongly controlled by structural features, like the ring-faultsystem at the borders of the caldera collapse. The use of detailed volcanological, mathematicaland statistical procedures also make it possible to obtain a detailed picture of eruptive hazardsin the whole Neapolitan area. The complex behaviour of this caldera, involving interactionbetween magmatic and geothermal phenomena, sheds light on the dynamics of the mostdangerous types of volcanoes in the world.

The Campi Flegrei volcanic field (Fig. 1) isthe largest feature of the Phlegraean VolcanicDistrict, which includes the islands of Procidaand Ischia, as well as submarine vents in thenorthwestern Gulf of Naples. It lies within theCampanian Plain, an 80x30-km-wide graben,formed since the Pliocene and still affected bycontinuous subsidence at a rate of 1.5–2mm a−1

(Dvorak & Mastrolorenzo 1991). The most con-spicuous geological feature of Campi Flegrei isa 12-km-wide collapse caldera outlined, on land,by a discontinuous ring of hilly morphology withinward-facing scarps enclosing the volcanic field.The caldera rim, inferred in scattered pointsby the stratigraphic sequence (Camaldoli Hill),ranges in elevation between about 450 m anda few tens of metres, and continues for aboutone-third of its depth below sea-level, formingthe Bay of Pozzuoli. A few marine volcanic relictsrepresent the submerged part of the volcanicfield. The southern limits of the caldera (belowthe sea) are poorly known and mainly inferredby geophysical investigations. A near-verticalfault system defines the geometry of the centraldepression which characterizes the nearlycircular caldera structure.

The intracalderic area of Campi Flegreishows the typical features of a volcanic fieldcharacterized by different landforms (Fig. 1). It

includes closely grouped volcanic hills, coalescedcraters, depressions bordered by steep, erodedvolcano flanks and fault scarps, crater-fillinglakes, and relicts of ancient marine terraces. Thefew major intracalderic plains of Fuorigrotta,Soccavo, Pianura, S. Vito, Quarto and LaSchiana are associated with localized subsidenceareas due to local movement following eruptiveepisodes. The coasts of Campi Flegrei can begrouped in two principal types: beaches at thetermination of alluvial plains, and the steep,mostly tuffaceous, cliffs – many of which arenatural sections of volcanic cones eroded by thesea or dislocated by faults (volcano-tectonicevents). Littoral dunes and lagoons (Patria,Fusaro, Miseno and Lucrino lakes) created bysand-bars complete the beautiful physical settingof the territory.

Accurate geochronological analyses nowsupply an objective control on stratigraphiccorrelations and more effective constraints onthe ages of the main volcano-tectonic events.The Campanian Ignimbrite (VEI=6), whichhas been often thought to be associated with theearly caldera depression (e.g. Rosi & Sbrana1987) was erupted about 39 000 a BP (De Vivoet al. 2001), while the subsequent largest event,which generated the Neapolitan Yellow Tuff(NYT, VEI=5), occurred c. 15 000 a BP (Deino

26 G. DE NATALE ET AL.

et al. 2004), and caused the inner part of thecaldera to collapse by several hundred metres.Indeed, the inner caldera rims, recognizable inthe field and via geophysical methods, coincidewith a deep dislocation in the NYT formation.Since the NYT eruption, volcanic activity havebeen restricted to smaller events within thecaldera. To date, at least 60 eruptions have beenrecognized, from some tens of scattered vents(monogenetic events).

The caldera area is inhabited by about 1.5million people. Its magmatic system is still active,as testified by the 1538 AD Monte Nuovo erup-tion (Di Vito et al. 1999); the recent bradyseismicepisodes in 1969–1972 and 1982–1984 that havegenerated a net uplift of 3.5 m around the townof Pozzuoli; and the widespread occurrence offumaroles and thermal springs. The combinationof dense urbanization and very active short-termdeformation make the volcanic risk in the areavery high.

This work synthesizes the main features ofrecent unrest at Campi Flegrei caldera to presenta coherent model for unrest episodes and togain insights into the main issues of volcanichazard, pre-eruptive scenarios and the impact ofpyroclastic flows and falls.

Ground movements and recent episodes ofunrest

Studies of slow movements at Campi Flegreibegan with observations of sea-level markerson Roman coastal ruins, which were sensitiveto large, secular deformation. The first – andmost studied – archaeological site was theRoman market-place (Macellum) in Pozzuoli.The monument has been the object of intensiveresearch, starting shortly after its excavationin 1750 (Breislak 1792; Forbes 1829; Niccolini1839, 1845; Babbage 1847; Lyell 1872; Gunther1903; Parascandola 1947). A marine mollusc,Lithodomus lithophagus, has bored into and leftshells in three 13-m standing marble columns,recording the vertical movement of the groundwith respect to sea-level. Parascandola (1947)presented the first reconstruction of historicalground movements, later modified by Dvorak &Mastrolorenzo (1991), and, in the last few years,by Morhange et al. (1999) (Fig. 2).

The Pozzuoli area has been characterized,since at least Roman times, by subsidence at arate of about 1.1–1.7 cm a−1, interrupted threetimes, once in the Middle Ages, and once about40 years before the last eruption in 1538. After

Fig. 1. Schematic map of Campi Flegrei caldera showing the main structural features (redrawn fromDe Natale et al. 1991).

27CAMPI FLEGREI CALDERA DYNAMICS

the 1538 eruption, the subsidence continued atthe previous rate, until the end of 1960s, whenrapid uplift started again, totalling about 1.7 mbetween 1969 and 1972. There followed 10 yearsof virtually no movement (a net subsidence ofless than 30 cm) until 1982, when renewed upliftraised Pozzuoli by almost 2 m up until the endof 1984 (Fig. 2). The last two episodes of unrest(called bradyseisms, from the Greek term for‘slow earthquake’) aroused concern about a pos-sible impending eruption, causing the authoritiesto evacuate the town of Pozzuoli. As a result ofthis concern, and also beacuase of the risk ofdamage to buildings caused by earthquakes andstatic deformation, Pozzuoli was evacuated.

The pattern of ground deformation during therecent bradyseismic uplift resembled an almostcircular lens centred near Pozzuoli. Much recentresearch, expecially since the 1982–1984 unrest,has focused on the strongly confined nature ofthe deformation (decaying radially by more than95% over 3 km from the maximum), on the largescale of the deformation itself (up to 1.8 m in twoyears, as measured at Pozzuoli harbour) and onthe fact that the deformation did not culminatein an eruption. Thus, although the concentrationof the deformed area indicated can be explainedby the very shallow depth of the magma chamber(see, for instance, Berrino et al. 1984), the largeamount of uplift, interpreted in terms of theclassical point-source model (Mogi 1958) wouldacquire overpressures in the source exceedingsome hundreds of MPa, depending on the sourceradius, large enough to be sustained by thestrength of the host crust (De Natale et al. 1991).To accommodate more reasonable overpressures(roughly equivalent to the crust’s tensile strength,about 10 MPa or less) numerous models haveproposed a range of crustal rheologies, fromheterogeneous elastic models (e.g. Bianchi et al.1987), to visco-elastic or elasto-plastical. Regard-ing the source geometry, since the markedsymmetry of the vertical deformation suggestedan axial-symmetrical source, generally it hasbeen interpreted either as a circular source (e.g.Berrino et al. 1984; Bonafede et al. 1986) or asan oblate spheroid (e.g. Bianchi et al. 1987; DeNatale et al. 1997), eventually degenerating intoa horizontal sill (e.g. Dvorak & Berrino 1991).The oblate shape was generally preferred becauseit maximizes the amount of vertical displace-ment, thus helping to minimize the extremelylarge overpressures required to model about 2 mof uplift. Visco-elasticity has been used both asa diffuse rheology of the whole medium (e.g.Bonafede et al. 1986), and as localized around themagma chamber due to very high temperatures(Dragoni & Magnanensi 1989). De Natale et al.(1991), however, showed that the large amountof uplift recovered during the subsequent sub-sidence implied that the elastic response musthave continued for at least five to ten years afteruplift has ceased. Any viscous effect, therefore,must have been small over the same time intervaland, hence, negligible during the two years ofuplift itself. De Natale & Pingue (1993) firstshowed that the concentration of ground defor-mation, whose shape remained unchanged overall the episodes of rapid uplift and subsidence,could instead be explained by localized plasticitymarking the ring-faults at the caldera borders.Passive slip along these faults during inflationwould produce a constant and concentrated

Fig. 2. (a) Schematic history of vertical movements atMacellum in Pozzuoli, known as Serapis Temple(after Bellucci et al., 2006, paper 8 of this volume,pp. 141–157). Black circles represent the constraintsfound from radiocarbon and achaeologicalmeasurements by Morhange et al. (1999); white circles(post-1538) represent inferences by Dvorak &Mastrolorenzo (1991); (b) Vertical grounddisplacements as recorded at Pozzuoli harbour bylevelling data in the period 1969–2005 (Macedonio &Tammaro 2005; Del Gaudio et al. 2005).


pattern of deformation, almost independent ofthe source depth (see also De Natale et al. 1997).

Very recently, Trasatti et al. (2005) presenteda model in which the concentration of grounddeformation is explained by the elasto-plasticbehaviour of the whole crust, with a lower plasticthreshold within the caldera and, consequently,a lower effective rigidity of the inner calderastructure. However, Troise et al. (2004) havedemonstrated that the presence of ring-faults atthe inner caldera borders is much more effectivein producing a concentrated deformation withrespect to any realistic heterogeneity of rigidity.They also pointed out that the effect of ring-faults on secular subsidence should produce alarger deformed area with respect to thatinvolved in fast uplift and subsequent partialrecovery episodes.

Another striking feature of the recent unrestat Campi Flegrei was the occurrence of a certainamount of subsidence following the upliftepisodes (about 80 cm between 1985 and 2005).Oscillation of ground level between uplift andsubsidence episodes, common to other similarcalderas like Yellowstone (Arnet et al. 1997) isdifficult to interpret in the light of classicalmodels of inflation and deflation of magmachambers, when no eruption (with consequentdeflation) occurs at the end of uplift episodes.In addition, the fast subsidence phase followingthe 1984 uplift has been punctuated by somesmall uplifts, of maximum amounts between1.5 and 11 cm, with an almost constant periodof five years (Gaeta et al. 2003). Each of thesmall uplifts was also accompanied by micro-earthquake swarms lasting from days to weeks(Bianco et al. 2004). The occurrence of earth-quakes also characterized the two unrest episodesof 1969–1972 and 1982–1984, whereas no localseismicity has been observed in the absence ofuplift. Weaker seismic activity was recordedduring the 1969–1972 unrest, whereas a largenumber of earthquakes (about 15 000), withmagnitudes ranging from about zero to five,was recorded during 1982–1984 (De Natale &Zollo, 1986; De Natale et al. 1995). A magnitude4.2 earthquake with an epicentre in Pozzuolioccurred on 4 October, 1983.

Figure 4 synthesizes the main features charac-terizing the unrest in 1982–1984. Most of theseismicity occurred in the period March–April1984, including a swarm of 550 earthquakes on1 April. Figure 3 shows the most representativepicture of seismicity during the 1982–1984 unrestin terms of earthquake density maps, computedusing the Bayesian method of Presti et al. (2004)for the 450 best-recorded events during theperiod. Troise et al. (1997, 2003) gave a coherentmodel for earthquake occurrence during unrestepisodes, in terms of static stress changes due

Fig. 3. (a) Earthquake density in the x–y plane for 370events recorded in the Campi Flegrei area between1973 and 1984. Events were located using a minimumnumber of eight arrival-time readings (after Prestiet al. 2004); (b) and (bppppp ) show, respectively, marginalearthquake densities along west–east and north–southsections centred at the crater axis (after Presti et al.2004).


to the combined effect of the source of uplift,regional stress and the presence of ring-faults.A schematic picture of the model for earthquakeoccurrence during uplift episodes is shown inFigure 5b. Earthquake locations are focusedin zones of maximum Coulomb fracture stresschange (in red). Expected fault mechanisms forthis model also agree with observations, havingnormal faulting dip components.

The Campi Flegrei substructure and theproblem of magma-chamber location

The Campi Flegrei substructure can be con-strained by several types of datasets. The firstcomes from deep drillings made by the Italian

Oil Agency (AGIP 1987) in the 1970s. Theymeasured temperatures, porosity and density inseveral wells in the Mofete area, on the eastern-most caldera rims, and at some wells at San Vito,about 2 km north of Pozzuoli (Fig. 1). In addi-tion, information on seismic velocities has beenobtained from local earthquakes recorded bythe permanent and temporary seismic networksoperating in the area. Most of earthquake datawere collected during the 1982–1984 unrest, bythe permanent monitoring network as well asa temporary digital network by the Universityof Wisconsin (UW) in 1984 (Aster et al. 1992).Data from the temporary network were usedto obtain the first VP (for P-wave velocity) andVS (for S-wave velocity) tomographic image of

Fig. 4. Map of geophysical observations at Campi Flegrei during the 1982–1984 unrest episode. Contours ofvertical displacements (in cm) and earthquake epicentres are shown; the projection of the collapsed zone asmodelled from gravity anomalies is superimposed on the depth section of the hypocentres. Composite focalmechanisms computed for the different seismic zones are indicated (De Natale et al. 1995). Also shown is thelocation of the possible magma chamber, as inferred by Ferrucci et al. (1992) (after De Natale et al. 2001).


Campi Flegrei (Aster & Meyer 1987). This imagehas recently been modified by inversion of animproved earthquake dataset, containing about400 earthquakes recorded in the period 1983–1984 by both the permanent and the UW tempo-rary network (Vanorio et al. 2005; Vinciguerraet al. 2006).

The third, and most informative, source ofdata about the shallow structure of CampiFlegrei and the surrounding area is the activeseismic tomography survey made by sea shotsin the Gulf of Pozzuoli (Zollo et al. 2003). Fromthis dataset, profiles of seismic velocity havebeen obtained to depths of about 5 km, witha generally high resolution due to the highseismic frequencies involved and the design ofthe experiment. In addition, a conversion fromP- to S-waves has been recorded during the 1987active seismic experiments in the Gulf of Pozzuoli(Ferrucci et al. 1990), and during regional/teleseismic events (Ferrucci et al. 1992; De Nataleet al. 2001).

Finally, useful constraints on seismic veloci-ties can be inferred by laboratory experimentson Campi Flegrei rocks, aimed at determiningseismic velocities under realistic conditions onrock samples (Zamora et al. 2001; Vanorio et al.2005; Vinciguerra et al. 2006).

Further constraints are also available fromgravimetric Bouguer anomalies determined byAGIP (1987) over a dense grid covering theCampi Flegrei area, and from petrological esti-mates of magma residence depths (Rosi &Sbrana 1987; Pappalardo et al. 2002).

Primary evidence for the shallow structureof Campi Flegrei is the presence of an almostcollapsed area, with a radius of about 2–3 km,around Pozzuoli. This feature, which is wellconstrained by gravimetric Bouguer anomalies(Rosi & Sbrana, 1987) and by recent activetomographic data (Zollo et al. 2003), is inter-preted as the inner caldera formed 15 000 a BPby the so-called Neapolitan Yellow Tuff erup-tion. With a maximum depth of 3 km, the col-lapsed basin has been filled by low-density andweakly coherent material erupted after calderaformation.

Figure 6a & b shows the maps of VP andVP/VS velocities as determined from local earth-quake tomography in this paper. These havebeen obtained by inverting 2500 P- and 1257S-wave arrival times from 331 earthquakesdetected by seismic stations (of both permanentand temporary networks) operating in the areaduring 1973–1984. Earthquakes were selectedto have a minimum number of eight arrival-time recordings. Data were inverted usingthe SIMULPS13Q program (Eberhart-Philips1998; Thurber 1993) with finite-difference 3D

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Fig. 5. (a) Maximum Coulomb stress changes for aspherical source (radius 1 km, depth 4.5 km andDP 5 MPa) embedded in elastic half-spaces and inthe presence of a background extensional stress sR

amounting to 10 MPa (after Troise et al. 2003);(b) schematic view of the movements along the ringfaults produced by a spherical pressure source,compared with those generated by anormal-fault local earthquake (after De Natale et al.2001). The apparent thrust fault movement along thering faults is not seen in earthquake data, so it must beaseismic (see Troise et al. 1997).


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ray-tracing on a 1-km grid. Levemberg–Marquardt dampings as lP=7, lS=12 were cho-sen to optimize the trade-off between decrease ofresiduals and variance norm. The starting modelwas a half-space with VP=3.0 km s−1 and VP/VS=1.8, which minimized the travel time residu-als. After four inversion steps, the RMS reduc-tion was 40%. The results indicate a zone, locatedaround Pozzuoli between 0 and 3 km depth, withextremely high Vp /VS (higher than 2.0), consis-tent with previous results (Aster & Meyer 1988;Vanorio et al. 2005) and average values for VP

(2.9 km s−1) higher than those determined inlaboratory on dry tuff samples, but agreeing wellwith laboratory measurements of saturated tuffs(Vinciguerra et al. 2006). The implication is thatPozzuoli overlies a system of shallow aquifersin an almost cylindrical volume, about 1 km indiameter, of water-saturated and fractured rock.The maximum vertical extent of the aquiferscan be constrained from pressure–temperatureconditions found in the Mofete wells. These indi-cate critical water-vapour equilibrium at a depthof about 3 km, so that water cannot be liquidbelow this depth threshold. It is also important tonote that temperature profiles at Campi Flegreiwells give maximum temperatures of about450 °C at 3 km, and respective average gradientsof about 150 °C km−1.

Local seismicity abruptly terminates at depthsof 3.5–4.0 km, suggesting a sharp transition frombrittle to ductile behaviour (see Hill 1992).Indeed, conversion of P- to S-waves from activeshots in the Gulf of Pozzuoli and regionalseismicity (Ferrucci et al. 1992), as well fromteleseismic waves in the inner caldera (De Nataleet al. 2001), suggest the presence of magma or, atleast, hot ductile rock, at depths of 3.5–4.0 km.However, seismic tomographic surveys have notfound the low-velocity anomalies expected atthese depths if magma were present. Instead, theyindicate zones of high velocity (6.0–6.2 km s−1)between 4 and 5 km below the centre of thecaldera.

In principle, if the observed P/S conversionsare generated at the surface of a small magmasill, the sill could go undetected by transmissiontomography, which at depths of 4–5 km has aresolution hardly better than 1 km3. However,the problem of magma detection by tomographicmethods, mainly at calderas, is much moregeneral. Calderas, in fact, are volcanic areasformed by huge quantities of erupted magma,and hence imply the existence of large magmachambers, with typical volumes in the order of100–1000 km3 and more. Generally, although theuse of seismic tomography is now rather wide-spread worldwide, almost no evidence is foundfor large, shallow magma chambers at calderas,

compatible with petrologically inferred volumes.On the contrary, high velocities are more oftenfound at depths where shallow magma chamberswould be expected (Zollo et al. 1996; Chiarabbaet al. 1997; Chiarabba et al. 2000; De Natale et al.2004). At Long Valley caldera, for instance, shal-low bodies (marked by low VP) have been found(Weiland et al. 1995), but much smaller thanthe huge volumes of original magma chamber asrequired by petrological inferences. Such world-wide negative evidence for molten residualmagma could be interpreted as being due to thefact that residual magma after large eruptionsis generally strongly crystallized (i.e. a mush). Aneffective mechanism for sudden crystallizationwell before the long times required by coolingof considerable volumes of molten rock, is rapiddegassing during eruptions (De Natale et al.2004). In this hypothesis, subsequent large erup-tions could be driven by inflow of new volatiles(mainly water) in the mushy magma chambers,remelting the previous magma, and/or by newamounts of molten magma.

Zollo et al. (2003) interpreted the high-velocity layers below 4 km as limestones andinferred that any major magma reservoir shouldlie at greater depths. An alternative interpre-tation is that the high-velocity layers representsolidified magma. According to Schon (2004),seismic velocities, in the order of 6.0–6.2 km s−1 ata depth of 4 km are consistent both with solidi-fied lavas at 4–6 km and water-saturated lime-stones, but are too high for dry limestone. Giventhe maximum depth of 3 km for water to exist(from Mofete data), the interpretation could bethat the high-velocity zone reflects levels ofsolidified magma. Moreover, such an interpreta-tion naturally accounts for the presence of highvelocities in the depth range where residualmagma is expected after caldera formation(Burov & Guillou-Frottier 1999). The abrupttermination of seismicity below 4 km could alsobe explained in terms of brittle–ductile transitionat this depth, coinciding with an increase intemperature above 500–600 °C (Waite & Smith2002). Indeed, simple extrapolation of the tem-peratures and gradients measured at the deepestMofete well indicates minimum values of 600 °Cat depths of 4 km (AGIP 1987; Gaeta et al. 1998).A similar conclusion was reached by De Lorenzoet al. (2001), who analysed the anelastic attenua-tion of seismic waves to constrain a thermalmodel for the Campi Flegrei substructure.

In conclusion, provided that the subsurfaceconditions measured at the Mofete wells canbe considered representative of all the inner-caldera rocks, the shallowest magma chamberis most likely to be located at depths greaterthan 4–6 km. Most of the magma is expected


to be degassed and, hence, mostly crystallized,although small batches of molten magma mayexist and pass undetected by transmission tomo-graphy, especially if stored in small sills. Thestrong P to S conversions observed at depthsof about 3.5–4 km could be produced by such asill or, alternatively, by a lens of magmatic fluidslocated just above the magma chamber, whichis typical of any volcanic geothermal systems(Fournier 1999). Petrological data indicate pre-ferred depths for magma residence of 5–8 km and10–15 km (Pappalardo et al. 1999; Pappalardoet al. 2002). Such depths are very similar to

those inferred for magma storage below Vesuvius(Pappalardo et al. 2004; De Natale et al. 2006).The resulting tentative reconstruction of theCampi Flegrei substructure is shown in Figure 7.

A mixed magmatic fluid-dynamic modelfor unrest episodes at Campi Flegrei andsimilar calderas

The peculiar time behaviour of ground deforma-tion at Campi Flegrei (Fig. 2) with fast uplift andsubsequent subsidence without eruptions, calls

Fig. 7. Tentative schematic reconstruction of the Campi Flegrei substructure. The central aquifer is schematicallyshown in blue, as inferred from VP/VS anomalies. A possible shallow, small magma chamber could be located atc. 4–4.5 km depth. The schematic picture of the hypothesized magma chambers below 4–5 km is mainly based onpetrological data (e.g. Mastrolorenzo et al. 2004). The large residual magma chamber responsible for the caldera-forming eruptions could be now almost totally crystallized (mush), giving the high velocities inferred by recentactive tomography (Zollo et al. 2003). Below 10 km, petrological data infer another magma source, which couldbe in relation to the large sill under the Campanian volcanoes, as inferred by Judenherc & Zollo (2004). Deepermagma roots are based on regional tomographic studies by De Gori et al. (2001). The inner-caldera collapse,surrounded by buried ring-faults, is shown by continuous lines. The apparent limits of a larger caldera are shownby broken lines.


for a mechanism of inflow and outflow of fluidfrom the system. Several papers have interpretedsuch evidence in terms of the involvement of theshallow geothermal system in the unrest episodes(Bonafede 1991; De Natale et al. 1991; Gaetaet al. 1998; Troise et al. 2001; Chiodini et al.2003).

To test this interpretation we here considerthe ground deformation and microgravitychanges measured at this area (Battaglia et al.2006). Since ground deformation constrainsthe volume changes at depth and microgravityconstrains the mass change, the two datasetstogether can constrain the density of the fluidinvolved. In our analyses we firstly consider ahomogeneous, non-fractured medium; further,we show the modifications introduced by con-sidering the presence of ring-faults marking thecaldera borders.

Joint inversion of deformation and gravitychanges

The temporal gravity change Dg determined bydifferencing repeated gravity measurements iscontrolled by the sum of four elements: (1) thefree-air effect, proportional to the uplift h; (2)the water-table effect, Dgw; proportional to theeffective porosity ϕe and water-table changeδz; (3) the deformation effect DgD; and (4) theresidual gravity DgR, which depends on the masschange accompanying the deformation (Eggers1987; Battaglia et al. 2003).

We use the theoretical value of the free-airgradient rather than the local observed value(Berrino et al. 1984) to compute the free-airgravity correction, because the theoretical value

better represents the change of gravity due touplift at the caldera scale. Measured free-airgradients are strongly influenced by terrain andlocal effects that will cancel out in differentialgravity surveys (Lafeher 1991).

Since there are no direct measurements ofwater-table changes that cover the time span ofour study, we computed the water-table rechargeϕeδz using a cumulative rainfall departure (CRD)method (Xu & van Tonder 2001). In coastal aqui-fers, taking into account urban development,the recharge is about 30 to 40% of the CRD(Appleyard 1995). Because of the limited re-charge, the water-table effect is negligible duringthe time span of our investigation.

The deformation effect ∆gD is the result ofsubsurface mass redistribution due to a dilata-tional source (Walsh & Rice 1979). This effectis exactly zero if the source has a spherical sym-metry (Walsh & Rice 1979), but can stronglybias the density estimate for a horizontal penny-shaped crack (Table 1).

Our data include levelling, EDM and gravitymeasurements collected by the OsservatorioVesuviano (Fig. 8). The analysed periods havebeen chosen in order to include horizontal defor-mation data in time intervals comparable withthe other data, given that EDM data were onlyoccasionally collected in the considered period.First-order optical levelling data, referenced tothe first Benchmark in Naples, are available forthe whole period 1980–1995. Distance changesin an EDM network covering the caldera weredetermined for 13 baselines between 1980 and1983 (during inflation; Dvorak & Berrino 1991),and 21 baselines between 1991 and 1995 (duringdeflation; Berrino 1995). Our computations inthis paper will assume a homogeneous, elastic,

Table 1. Summary of the joint inversion of geodetic and gravity data*

Model χ2ν Z ∆V ∆P/µ ρ ρ* R A

1980–1983 (During inflation)

Poi 17 3.0 33 – 3.0 3.0 – –Sph a 16 2.8 28 0.009 3.1 3.1 1 1Sph b 16 2.8 28 0.07 3.2 3.2 0.5 1Pen 2 2.6 21 6x10−4 0.6 4.6 2.4 –

1990– 1994 (During deflation)

Pen 1.5 3.6 −5 −6x10−4 −1.1 2.3 1.6 –Poi 1.5 3.0 −6 − 1.4 1.4 – –Sph a 1.3 2.1 −6 −0.008 1.0 1.0 1.0 0.45Sph b 1.3 2.0 −5 −0.05 1.0 1.0 0.5 0.5

*χ2ν: chi-square per degree of freedom; Z: depth (km); DV: volume change (10−3 km3); DP: pressure change;

m: shear modulus; r: density – taking into account the deformation effect (g cm−3); r*: density – assuming thedeformation effect is negligible (g cm−3); R: radius (km); A: aspect ratio. Pen: penny-shaped crack; Poi: pointsource; Sph: prolate spheroid.


non-faulted medium; later in the paper we willdiscuss the modifications introduced in themodel, considering the presence of ring-faultsfirstly shown by De Natale & Pingue (1993) andDe Natale et al. (1997). The possible effect ofnon-elastic (i.e. visco-elastic) rheology is not con-sidered here but, as demonstrated by De Nataleet al. (1991), it should not be significant. Six grav-ity stations are available for our analysis between1981 and 1983, increasing to 10 stations between1991 and 1995 (Berrino 1994; Gottsmann et al.2003).

The data have been inverted using a weightedleast-squares algorithm with an exhaustive gridsearch. The covariance matrix of the solution,besides the measurement error, includes an errorof 2 mm, taking into account the instability ofthe measuring stations. The best-fit model isdetermined using a reduced chi-square test,and 95% confidence limits are computed by abootstrap percentile method (Efron & Tibshirani1986). We test three source geometries: a spheri-cal source (McTigue 1987), a vertical prolatespheroid (Clark et al. 1986; Yang et al. 1988), and

a horizontal penny-shaped source (Fialko et al.2001), all in an elastic, homogeneous, isotropichalf-space.

Our results are synthesized in Table 1 andFigure 9. They show that the deformation andgravity data during inflation are best describedby changes in a horizontal penny-shaped crack,between 3.2 and 5.4 km across and 2.5–3.5 kmbelow Pozzuoli. This implies that the increasein volume is between 0.021 and 0.027 km3 for afluid with density between 142 and 1115 kg m−3

(Fig. 9a). In contrast, the best-fit source forthe deflation period is a vertical prolate spheroid1.9 to 2.2 km deep, with aspect ratio (the ratiobetween the semi-minor and the semi-majoraxes) from 0.39 to 0.54; volume decrease from0.005 to 0.006 km3; and a density between902 and 1015 kg m−3

(Fig. 9b) The inversion isinsensitive to the magnitude of the spheroidsemi-major axis (Table 1).

The best-fit results from joint inversion ofground deformations and gravity changes sug-gest that the bradyseismic movements since1982 have been controlled by pressure changesin two water-rich horizons. The deeper horizon isa sill-like reservoir above the magma chamber,at 3.5–4 km depth. The upper horizon, at 1.8–2.4 km depth, has the form of a prolate spheroidand so has the depth and shape consistent withthe shallow geothermal systems from seismictomography studies and geothermal explorationwells (De Natale et al. 2001). The large differencein the source shape between the uplift (a penny-shaped crack) and subsidence (a vertical prolatesspheroid) phases is unambiguously constrainedby the large differences in the ratios betweenhorizontal (ur) and vertical (uz) deformations.This ratio is larger for vertically elongatedsources, whereas it is minimal for horizontallyelongated ones. In terms of maximum displace-ments (in this work we use line elongations ashorizontal data), a Mogi (1958) model has a ratioof about 0.35, which is also similar for sphericalsources in general. The following is an examplefrom our data: the ratio is 0.5 along the S–CEDM baseline (ur=0.4 m and uz=0.8 m at ben-chmark 25) between 1980 and 1983 (uplift); theratio is 0.85 along the ACA–BAC EDM baseline(ur=0.128 m and uz=0.151 m at benchmark 25)between 1990 and 1995 (subsidence).

To account for the observed patterns ofground deformation and gravity change, we pro-pose that the 1982–1984 inflation was driven byhypersaline brine and gas, which, escaping fromcrystallizing magma, accumulates in a horizontallens above the magma chamber. A self-sealedzone of impermeable material separates this lensfrom the shallow aquifers. When a major breachof the self-sealed zone occurs (some time between

Fig. 8. Sketch map of the levelling lines, EDM, gravitybenchmarks and source location at different periods oftime, as analysed in this work. (a) Between 1980 and1983 (during inflation); (b) between 1990 and 1995(during deflation).

36 G. DE NATALE ET AL.(a

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1983 and 1985), brine and gases migrate to thelower pressure and colder aquifers in the brittlecrust (Fournier 1999). The consequent fluidaccumulation in the shallower aquifer, wouldamplify the deformation, because the amount ofuplift increases as the depth of the pressuresource decreases. Subsidence starts when theresulting increase in fluid pressure and tempera-ture within the brittle aquifer leads to faultingand fracturing, which increases the permeabilityand allows an increase in the rate of discharge ofhydrothermal fluids. Lateral migration of fluidsoutside the caldera can be an efficient dischargemechanism (De Natale et al. 2001).

Although this model can explain bradyseimicmovements in terms of pressure changes andfluid migration in aquifers, it does not directlyaddress the process that initially disturbed con-ditions in the lower aquifer. One possibility isa sudden increased output from the magmachamber, perhaps as a response to new magmainjected from below. A magmatic trigger forthe unrest cannot be completely excluded by theprevious results obtained for a homogeneouscrust. Indeed, the effect of ring-faults at calderasmakes surface deformations almost insensitive tothe source depth, within a certain depth interval.Given that the source depth is only constrainedby surface displacement data, the use of a homo-geneous crust may underestimate its actual value.Thus, a pressure source with a centre at 5 kmdepth could optimally fit displacement data, withalmost the same overpressure, provided that theeffect of ring-faults is considered (De Natale et al.1997, 2001; Beauducel et al. 2004). If the depth ofthe sill source is increased to 5 km, the surfacegravity change caused by the intruded masswill be decreased, because this depends on 1/R2

(where R is the depth). Hence, a higher density ofintruding fluid would be required to produce theobserved gravity change. As a first approxima-tion (neglecting the change of deformation effect∆gD with depth), the decrease in gravity changefrom a depth of 3 to 5 km is in the order of 9/25,so that fluid density must be multiplied by 25/9to reproduce the observed changes. The resultingvalues would, therefore, lie in the range 394–3090 kg m−3, i.e. embracing the densities for bothwater and magma.

It is important to stress that the presenceof ring-faults would have almost no effect onthe depth of the deflation source, because theobserved high ratio between horizontal and verti-cal deformation can only be reproduced by avery shallow, vertically elongated source. In fact,the effect of ring-faults does not significally affect

the maximum horizontal/vertical displacementratio (Fig. 10), high values for which can onlybe justified by a source depth and shape similarto that found in a homogeneous crust. To provethis in a clear way, we have computed, andshowed in Figure 10, the maximum horizontal/vertical displacement ratio as a function ofsource depth and ellipticity (ratio between themaximum and minimum axes of a prolatespheroid, taken as 0.5, 0 and 2.0 respectively) inan axial-symmetrical homogeneous half-spacewith ring-faults. The resulting source depth andshape have values very similar to those found in acontinuous crust.

The subsidence phase following fast upliftat Campi Flegrei is thus most likely to be causedby a shallow source, elongated vertically. Thefast uplift itself, however, may be considered tohave been caused by changes in a deeper magmabody or a water/gas-filled reservoir. The pre-ferred model involves a first phase of upliftproduced by a sill-like source filled by magmaticfluids or magma itself; magmatic fluids thenmigrate into the shallow aquifers, amplifyingprogressively the ground deformation whichdepends on 1/R2 (R=depth). After reaching themaximum pressurization of the whole aquifer,water tends to flow out of the caldera, causingthe subsidence phase immediately following theuplift.

Fig. 10. Maximum horizontal/vertical ratio v. sourcedepth in the case of three sources with differentgeometries (prolate ellipsoid, sphere, oblate ellipsoid)embedded in an elastic half-space with ring-faultslocated to simulate Campi Flegrei Caldera.


Implications of the model for possiblepre-eruptive processes

The model obtained in our study has fundamen-tal implications for forecasting eruptions atCampi Flegrei and similar calderas. Part of theuplift (in our case the deformation between 1982and 1984) is amplified by perturbations of thegeothermal system. Hence, if the triggering pro-cess is magma intrusion, it must take place at thebeginning of the crisis; subsequent subsidence,in turn, represents lateral outflow of water froma shallow aquifer. An immediate consequence ofthis model is that the maximum uplift is notdirectly associated with a maximum in magmaoverpressure, but instead to a delayed responseof the geothermal aquifer. The maximum hazardfor an eruption is therefore at the beginning ofunrest, in correspondence with the magmaticinput. In addition, if the initial input correspondsto the injection of new magma, the cumulativeeffect of several unrests would be to accumulatemagma overpressure in the chamber. Theeruption could then occur when, after a certainnumber of consecutive unrests episodes, cumula-tive magma pressure overcomes the thresholdof rock strength (Fig. 11). Such a model is alsoconsistent with conditions observed in 1994 atRabaul Caldera, New Guinea (Nairn et al. 1995)where eruption from two vents began abouttwelve hours after the start of weak precursory

signals (a seismic swarm of low-frequency earth-quakes). Significantly, the eruption occurredten years after the end of a more than a decade-long episode of geological unrest involving morethan two metres of uplift and intense seismicity,which may have brought the crust to a state ofunstability. The model in Figure 11 may thusdescribe a general pattern of behaviour forcalderas. If so, eruption forecasts should relymainly on the detection of a new magmaticphase, which presumably has much less effect onobservable data with respect to the spectacularground deformation and seismicity which char-acterize a mature unrest. The magmatic input,representing a deeper source, should have aminor effect on surface deformation. Also impor-tant, at least in Campi Flegrei, is that the seis-micity started about two months after grounddeformation, so that seismicity cannot be reliedupon to reveal the very first stages of unrest.Obtaining data that are more sensitive to aninitial magma input from depth may thus requirehigh-sensitivity borehole dilatometry or micro-gravimetry.

The key results, therefore, are that, against thebackground of significant geophysical anomaliesobserved during unrest that does not includeeruptions, true eruption precursors can be verysmall and difficult to detect, so requiring extremecare in data acquisition and analysis. Such con-straints provide useful insights into improvingthe monitoring network at Campi Flegrei, evenif it is already one of the most well-monitoredvolcanic areas in the world. A very importantstep in future research should then be the physi-cal characterization of the most probable trueprecursors of eruptions, which could involveeither very small initial deformation and/or theformation of an eruption pathway by progressiveearthquake fracturing (Kilburn & Sammonds2005).

Probabilistic estimation of eruptive productshazard

The key issue for a highly explosive volcanicarea like Campi Flegrei, an extremely populatedregion, is the accurate estimation of hazards fromeruptions, and their impact on urbanz areas. Inparticular, it is important to estimate the prob-ability of occurrence, at each point of the area, ofthe two main eruption products which constitutethe highest risk: pyroclastic currents and thefallout products (ash, pumice, etc.). Despite theprogress made in the last decade on the rigorousmodelling of pyroclastic flows generated byexplosive eruptions (i.e. Neri et al. 2003) these

Fig. 11. Schematic view of vertical displacements as afunction of magmatic source overpressure. In thissketch, for a fluid-dynamic model with magmaoverpressure as the initial input, most of the grounddeformation is caused by a delayed response fromaquifers. In this case, the highest eruptive hazardshould be at the beginning of the unrest (shadedintervals), not during the phase with the largest uplift.Eruption occurs when the rock strength (dashed line)is overcome by the cumulative overpressure built byseveral consecutive periods of unrest.


methods are also too cumbersome even for pow-erful parallel computers to make a probabilisticanalysis, which requires a huge number of simu-lations to reproduce and randomize all of thelikely possible kinds of eruptions. Moreover,finite difference fluid-dynamic models requirea large number of parameters to be specified,which are mostly unknown, and cannot handlea large number of grain sizes in the flow, as it isobserved in the field (continuous granulometry).This is the reason why, for hazard evaluation,models which approximate the flows in terms ofsimple rheological models have been widely usedin the last two decades (Sheridan 1979; McEwen& Malin 1989). Rheological models, which usethe most general Bingham rheology, also includea particular example of the Newtonian model,allow in practice to model the heavier part ofthe flow, which flows on the ground under thecombined effect of the initial acceleration and

the topography. Such an approximation is goodenough if the flow is gravity-driven, and, asshown in many analyses (see, for instance, Neriet al. 2003) holds in particular for flows withgrain sizes higher than about 10 mm, which is avery low threshold for most of granulometrieseffectively observed in pyroclastic flows atCampi Flegrei. In addition to pyroclastic cur-rents, another important source of hazards isrepresented by light material in plinian columnswhich falls on the ground, and whose accumula-tion can cause roofs to collapse. Several theore-tical methodologies exist, almost all based on thefirst models of Suzuki (1983) and Connor et al.(2001). Although plinian columns are not themost common explosive type in this area, thiskind of phenomenon has occurred in the pastand its hazard potential can be quantified. Thebasic method used by volcanic geologists to maphazard potential is based on the direct mapping

Fig. 12. Probabilistic hazard map from pyroclastic currents in the Campi Flegrei area, computed from thedistribution of eruptions in the last 5000 years. The value gives the yearly probability that each small area of asquare kilometre will be hit by a pyroclastic current (after Rossano et al. 2004).


of past eruptive products (Connor et al. 2001).Obviously, such a method is largely unsatisfac-tory because the past eruptive history, providedthat it can be well reconstructed, gives a veryuneven and inadequate sample of the whole spec-trum of possible eruptions, making meaninglessany statistical assessment of hazards that hasbeen constructed on such a basis. Mainly at alarge volcanic field as large as Campi Flegrei,in fact, the probabilistic reconstruction of thehazard potential requires us to consider the factthat any type eruption that in the past hasoccurred once at a certain site, in principle canoccur at any place in the caldera. Recently,Rossano et al. (2004) presented a rigorousvolcanological–statistical procedure to construc-tion hazard maps from records of pyroclasticcurrents. The resulting map for Campi Flegrei,giving the yearly probability of any site being hitby a pyroclastic current, is shown in Figure 12.

In addition, the basic methodology can beeasily generalized to create hazard maps forany type of eruptive product. In the followingsection, we show a reconstruction of fallout haz-ard in the area based on a similar probabilisticmethod.

Pyroclastic fallout

The transport of tephra in the atmosphere, dueto the processes of convection, diffusion and set-tling, may be described by the following well-known convection–diffusion–migration equation(Friedlander 1977), written here in conservationform:

∂∂

∂∂

∂∂

∂∂

∂∂

∂∂

∂∂

∂∂

∂∂

χ χ χ χ

χ χ

tux

vy

wz

xK

x yK

y zx y

+ + +

= + + KKz

Vx

z

∂∂

∂∂

χ

χ- ts

(1)

where x is the concentration of tephra (kg m−3),vw(u,v,w) is the wind velocity vector, Kx, Ky, Kz

are diffusion coefficients, and Vts is the terminalsettling velocity. Where a low volume concen-tration of tephra is hypothesized, v, Ki and Vts

depend only on (x,y,z,t) but not on x, so theequation is linear. Given an expression for suchcoefficients and suitable boundary and initialconditions, eq. (1) may be solved numericallyto give x=x (x,y,z,t). The mass distribution oftephra after all the particles have settled may thenbe calculated by the integration of vertical fluxesat the surface, at a time sufficiently longer thanthe settling time of the smallest particles that we

are interested in. However, such an approach iscomputationally expensive: in order to calculatethe 2D distribution of tephra on the ground fortF>ts, we have to compute a 3D field for all timest<tF. According to Suzuki (1983), we can neglectvertical diffusion and convection in the atmo-sphere with respect to horizontal diffusion andconvection because the former has a muchsmaller effect than the latter above the atmo-spheric boundary layer. We have:

∂∂

∂∂

∂∂

∂∂

∂∂

∂∂

∂∂

χ χ χ χ χt

ux

vy x

Kx y

Kyx y+ + = + (2)

Now x has the meaning of mass of accumulatedtephra for unit surface (kg m−2). The settling termis implicitly taken into account by consideringthe solution of eq. (2) at the time correspondingto the fall time of the particles. Diffusion isassumed to be isotropic (Kx=Ky=K), with windvelocity depending only on z, and the diffusioncoefficient is assumed to be constant in space andtime.

The method to solve equation (2), and toobtain a simulated 2D distribution on the groundof tephra thickness as a function of a set of erup-tion parameters, is described in Mastrolorenzoet al. (2006).

Hazard maps are generated by simulatingmany eruptions, grouped according to VEI (vol-canic eruption index) from 0 to 5, each with adistinctive probability according to the volcanichistory of Campi Flegrei.

The different values of eruption parametersfor each VEI are assumed with distinctive prob-abilities according to the inferred features ofpast eruptions, on the basis of the study of theirdeposits. Wind velocity and direction statisticsare taken from meteorological archives (seeMastrolorenzo et al. 2006, this volume).

For each eruption of probability p, the heighth of deposited tephra at each point (xm, ym) ofthe map is calculated from the grid (xi, yi) ofthe possible eruptive vents. For each (xm, ym), weassign h to one of m bins Ii covering all possiblevalues, and add the probability p to the corre-sponding bin. Repeating this process for all erup-tions, we obtain a probability histogram of theheight of deposited tephra for each point (xm, ym).This gives the complete statistical description ofthe random variable h, so now we can computethe probability of h being higher than a giventhreshold, the mean value of h and so on, deriv-ing the corresponding isomaps. The hazard mapfor tephra fall so computed is shown in Figure 13,which gives the yearly probability of a fall-outloading larger than 200 kg m−2 (h>0.2 m).


Conclusions

Campi Flegrei has been one of the world’s mostactive calderas in the last 40 years, generatingunrest phenomena not followed by eruptions,characterized by huge ground deformation andseismicity. Data from Campi Flegrei shed lighton several peculiar features of caldera dynamics,and help to constrain many details of the sub-structure. In particular, large amounts of uplift(up to 3.5 m in 15 years) and uplift rates (up to1 m a−1), sometimes followed by partial subsid-ence, are explained in terms of an initial magmainput followed by pressurization of the geother-mal system, which then progressively deflatesby water outflow. The water deflation phase,associated with the subsidence following largeuplift episodes, is strongly constrained by gravityand ground deformation data, which uniquelyindicate that the subsidence source is a shallow(about 2 km), vertically elongated type, withfluid characterized by densities in the range500–1 000 kg m−3.

The mixed magmatic–geothermal model forunrest episodes takes into account all the possibleeffects linked to the structural effects of calderason ground deformation. Such effects, firstlyevidenced by modern research on this caldera,mainly involve the passive slip of collapse ring-faults in response to the overpressure in magmachambers and aquifers. Local seismicity is shownto be also linked to the Coulomb stress changesproduced by overpressure sources and ring-fault

passive slip, in a background normal-faultingregional stress regime.

Studies of the Campi Flegrei substructure,involving gravimetric data, earthquake loca-tions, active and passive seismic tomography andlaboratory experiments on rock samples, allowus to make some inferences about the locationof the brittle–ductile transition, shallow aquifers,and the thermal–rheological state of rocks downto 5 km depth. Clear evidence for magmachambers only comes from petrological studies,indicating two main depth ranges for magmaaccumulation (5–8 and 11–15 km). As in mostof the caldera areas, no clear seismological evi-dence exists for large volumes of low P-velocities,thus indicating that the residual magma chambershould rather be characterized by a high degreeof crystallization, probably because of degassing.

While the models for the occurrence of unresthelp to constrain possible pre-eruptive scenarios,indicating the additional complexity of eruptionforecast at calderas, with respect to central volca-noes, the use of accurate modelling and statisticalmethodologies allow us to reconstruct, fromdetailed studies of past eruptive history, theprobability of occurrence of the main destructiveevents linked to pyroclastic product emission inthe whole Campi Flegrei and Neapolitan area.

The complex dynamics evidenced at CampiFlegrei, characterized by strict magma–waterinteraction not only during eruptions, and theproblems linked to location and rheology of

Fig. 13. Probabilistic hazard map derived from the tephra fallout pattern in the Campi Flegrei area (afterMastrolorenzo et al. in prep; see also Mastrolorenzo et al. 2006, this volume). The map gives the yearly probabil-ity that each 1 km2 area will experience a pyroclastic fallout loading of more than 200 kg m−2.


magma chambers, all indicate the need fordetailed, direct studies of caldera substructureby deep drilling. Detailed study of Campi Flegreicaldera, which represents a typical example ofpartially submerged super-volcano located ina densely populated urban area, thus gives aunique opportunity to deeply understand thedynamics of the most dangerous volcanoeson Earth, representing the most catastrophicnatural hazards, after giant meteoritic impacts.

We to thank C.R.J. Kilburn for useful suggestions andcomments, which greatly improved the quality of thepaper. Our research was partially supported by MIUR–FIRB 2001 funds (RBAU01M72W_001).

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