G 3. AN ELECTRONIC JOURNAL OF THE EARTH SCIENCES Published by AGU and the Geochemical Society

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1 Geosystems G 3 AN ELECTRONIC JOURNAL OF THE EARTH SCIENCES Published by AGU and the Geochemical Society Review Volume 6, Number 7 13 July 2005 Q07005, doi: ISSN: Relative roles of rifting tectonics and magma ascent processes: Inferences from geophysical, structural, volcanological, and geochemical data for the Neapolitan volcanic region (southern Italy) Monica Piochi, Pier Paolo Bruno, and Gianfilippo De Astis Osservatorio Vesuviano, Istituto Nazionale di Geofisica e Vulcanologia, Via Diocleziano 328, I Napoli, Italy (monky5@ov.ingv.it) [1] The Neapolitan volcanic region is located within the graben structure of the Campanian Plain (CP), which developed between the western sector of the Appenine Chain and the eastern margin of the Tyrrhenian Sea. Two volcanic areas, spaced less than 10 km apart, are situated within the CP: the Somma- Vesuvius Volcano (SVV) and the Phlegraean Volcanic District (PVD). SVV is a typical stratovolcano, whereas PVD, including Campi Flegrei, Procida, and Ischia, is composed mostly of monogenetic centers. This contrast is due to different magma supply systems: a widespread fissure-type system beneath the PVD and a central-type magma supply system for the SVV. Volcanological, geophysical, and geochemical data show that magma viscosity, magma supply rate, and depth of magma storage are comparable at PVD and SVV, whereas different structural arrangements characterize the two areas. On the basis of geophysical data and magma geochemistry, an oblique-extensional tectonic regime is proposed within the PVD, whereas in the SVV area a compressive stress regime dominates over extension. Geophysical data suggest that the area with the maximum deformation rate extends between the EW-running 41st parallel and the NE-running Magnaghi-Sebeto fault systems. The PVD extensional area is a consequence of the Tyrrhenian Sea opening and is decoupled from the surrounding areas (Roccamonfina and Somma-Vesuvius) which are still dominated by Adriatic slab dynamics. Spatially, we argue that the contribution of the asthenospheric wedge become much less important from W-NW to E-SE in the CP. The development of the two styles of volcanism in the CP reflects the different tectonic regimes acting in the area. Components: 15,031 words, 9 figures, 1 table. Keywords: volcanic styles; tectonic setting; Neapolitan volcanic region. Index Terms: 8109 Tectonophysics: Continental tectonics: extensional (0905); 8434 Volcanology: Magma migration and fragmentation; 8499 Volcanology: General or miscellaneous. Received 26 November 2004; Revised 19 April 2005; Accepted 27 April 2005; Published 13 July Piochi, M., P. P. Bruno, and G. De Astis (2005), Relative roles of rifting tectonics and magma ascent processes: Inferences from geophysical, structural, volcanological, and geochemical data for the Neapolitan volcanic region (southern Italy), Geochem. Geophys. Geosyst., 6, Q07005, doi:. 1. Introduction [2] Volcanic activity can generate composite (polygenetic) volcanoes or volcanic fields characterized by monogenetic vents depending (1) on the ability of feeder-dikes to coalesce, (2) on the progressive evolution and geometry of the associated magma chamber, and (3) on the stress field over the magma source region [see Cañon-Tapia and Walker, 2004; Gudmundsson, 1988, 1990, 1998; Marsh, 2000; Takada, 1989]. More specifically, the factors regulating the style of volcanism are Copyright 2005 by the American Geophysical Union 1 of 25

2 Figure 1. (a) Magmatic provinces and main tectonic lineaments in Italy (modified from Beccaluva et al. [1991] and Acocella et al. [1999]). The slab is the inferred front of the Ionian subducted plate. (b) Geological and tectonic sketch map of the Campanian Plain showing distribution of sedimentary sequences and PVD, SVV, Roccamonfina, Ventotene, and S. Stefano volcanic areas (modified from Orsi et al. [1996] and reprinted with permission from Elsevier). Sources of volcanological and structural data: Di Girolamo and Stanzione [1973]; Finetti and Morelli [1974]; Rosi and Sbrana [1987]; Beccaluva et al. [1991]; Orsi et al. [1996]; Acocella et al. [1999]. For further details, see text. (1) magma viscosity [Wylie et al., 1999], (2) rate of magma intrusion and extrusion [Hildreth, 1981], (3) rate of magma percolation [Hildreth, 1981], and (4) deformation rate in the lithosphere [Nakamura, 1986]. However, magma supply rate, tectonic setting, stress fields and mechanical state of the lithosphere have to be considered the most important [Cañon-Tapia and Walker, 2004; Fedotov, 1981; Gudmundsson, 1988, 2002; Takada, 1994a]. [3] The Neapolitan volcanic region (Figures 1a and 1b), one of the most hazardous volcanic areas in the world, consists of the Somma-Vesuvius Volcano (SVV) and the nearby Phlegraean Volcanic District (PVD). PVD, less than 10 km north of the SVV, includes the volcanic fields of Campi Flegrei, Procida and Ischia islands, and a number of submerged vents, aligned in a NE-SW direction. In the last decade, significant efforts have been made to better understand the volcanic and magmatic evolution of SVV and PVD [e.g., Arrighi et al., 2001; Ayuso et al., 1998; Civetta et al., 1991a, 2004; Civetta and Santacroce, 1992; Cortini and Hermes, 1981; D Antonio et al., 1999a, 1999b; De Astis et al., 2004; De Vivo et al., 1993; De Vivo and Rolandi, 2001; Di Vito et al., 1999; Marianelli et al., 1999; Marinoni, 2001; Orsi et al., 1995, 1996; Rolandi et al., 1998; Santacroce, 1987; Santacroce et al., 1993; Spera et al., 1998; Pappalardo et al., 1999, 2002; Piochi et al., 1999, 2004, 2005; Turi and Taylor, 1976]. In addition, seismic, gravity and aeromagnetic surveys have been carried out in order to define the main structural features of the Neapolitan area and to recognize submerged and buried tectonic lineaments both at regional and local scales [e.g., Azienda Generale Italiana Petroli (AGIP), 1981; Auger et al., 2001; Berrino et al., 1993; Bruno et al., 1998, 2000, 2002a, 2003; Ferrucci et al., 1989, 1992; Judenherc and Zollo, 2004; Rosi and Sbrana, 1987; Zollo et al., 2002, 2003]. Geochemical monitoring data has also contributed to the knowledge of the magmatic and structural setting of both PVD and SVV [e.g., Caliro et al., 1999; Chiodini et al., 2001a, 2001b, 2of25

3 2003; Frondini et al., 2004; Inguaggiato et al., 2000; Panichi et al., 1992]. [4] Despite this large number of studies, the mutual relationships among volcanism, tectonics, magma ascent and geodynamic processes have not been investigated. In this paper we want to address this question by merging the large amount of volcanological, geophysical and geochemical data available for the Neapolitan area and to develop a general understanding of the causes of polygenetic versus monogenetic volcanism in a postcollisional margin. PVD and SVV differ in important ways: (1) a high number of small monogenetic volcanoes occur in the Phlegraean area, whereas (2) a central-type (polygenetic) volcano occurs in the Vesuvian area. This distinction emphasizes the influence of regional tectonics on the style of volcanism and magma ascent. Our results contribute to the still open debate about the geodynamic evolution of Southern Italy and related magmatism [see Argnani and Savelli, 1999; Ayuso et al., 1998; Beccaluva et al., 1991; Gasperini et al., 2002; Peccerillo, 1999, 2001; Schiano et al., 2001; and references therein]. In particular, the geodynamic evolution of the Campanian Plain and its relationship to volcanic activity is for the first time discussed in this paper. Finally, our results suggest that information collected on the SVV and PVD magma supply systems, as well as knowledge of the mechanisms leading to changes in the supply geometry, can be a further diagnostic tool for volcanic hazard assessment. 2. Regional Geology [5] The Neapolitan Volcanic Region is part of the Italian Volcanic Belt and lies in the Campanian Plain (CP), between the western side of the Southern Apennine Chain and the eastern border of the Tyrrhenian abyssal plain (Figures 1a and 1b). Geological and geophysical evidence as well as compositional features of magmas erupted in south-central Italy suggest that subduction of a slab made of oceanic lithosphere from the relict Ionian basin and of thinned continental lithosphere of the Adriatic, Sicily and N-Africa areas, occurred beneath the Appennine [e.g., Lucente et al., 1999; Peccerillo, 1999; Savelli, 2001; Selvaggi and Amato, 1992; Serri et al., 1993]. During late Miocene-early Pliocene, the opening of the Tyrrhenian Sea (8 Ma to present [e.g., Scandone, 1979; Doglioni, 1991]) and the SE migration of the Calabrian arc followed the rollback (and possible detachment) of the subducted Ionian plate under Calabria [Gvirtzman and Nur, 2001; Piromallo and Morelli, 1997; Selvaggi and Chiarabba, 1995; Wortel and Spakman, 2000]. As a consequence of this tectonism extensional structures were produced and persistent volcanism occurred in Campania, as well as in neighboring Latium (Figure 1a) regions. Extension in the Tyrrhenian basin and compression in the Apenninic chain coexisted, with a progressive shifting of the rift/thrust belt/ foredeep system toward the Adriatic-Ionian foreland and development of extensional processes within the inactive thrust belt [Meletti et al., 2000, and references therein]. [6] As a result of the motions of the Tyrrhenian and Ionian blocks, the CP became a structural depression bordered by NW-SE and NE-SW trending faults [D Argenio et al., 1973; Hippolyte et al., 1994; Ippolito et al., 1975]. NW-SE fault system dominated the CP formation during Plio- Quaternary age, whereas the associated and coeval NE-SW transverse faults included normal and strike-slip components of motion [e.g., Acocella et al., 1999; Mariani and Prato, 1988]. [7] Radiometric ages available for the Campanian volcanoes, including also Ventotene and Roccamonfina (Figures 1a and 1b), indicate that volcanism started at about 1.5 Ma in the northernmost sectors of the region [Giannetti et al., 1979; Giannetti, 2001; Metrich et al., 1988] and at about 0.36 Ma in the SVV area [Brocchini et al., 2001; De Vivo et al., 2001]. Thus, during Plio-Pleistocene times, the volcanism progressively migrated from the Ventotene-Roccamonfina volcanoes ( Ma) to the Neapolitan area (360 ka to present time). Magmas generated in these centers show distinct enrichments in K 2 O and incompatible elements and allow one to distinguish two magmatic zones: the Ernici- Roccamonfina province and the Campania province, which includes the SVV and PVD [Peccerillo, 2002, and references therein]. On the basis of their geochemical and isotopic features, both these provinces are rather different from the northernmost Roman Province, which is characterized, on average, by higher 87 Sr/ 86 Sr and light ion lithophile elements/high field strength elements (LILE/HFSE) ratios [Beccaluva et al., 1991; Duschenes et al., 1986; Serri et al., 1993]. (LILE: Sr, K, Ba; HFSE: Th, Ta, Nb, Ce, P, Zr, Hf, Sm, Y, Yb.) Note also that the southern Latium and Campania regions are characterized 3of25

4 Figure 2. (a) Fault patterns in the CP obtained from active seismic surveys and from geological data (dotted lines represent inferred fault location). (b) Bouguer anomaly data in mgal (modified from Carrozzo et al. [1986]). Isolines are plotted every 4 mgal. The labeled areas indicate the basins. (c) Aeromagnetic map of the study area in nt (total field) [AGIP, 1981]. Isolines are plotted every 50 nt. (d) Heat flux distribution. Isoline spacing is 50 in mw/m 2. Note: color filling qualitatively shows maxima and minima of geophysical anomalies. by two deeply rooted structural features which probably separate two different mantle domains: (1) the NE-SW striking Ortona-Roccamonfina line [Billi et al., 1997; Mantovani et al., 1996; Patacca et al., 1990] and (2) the well-known WNW to E-W trending lithospheric discontinuity which corresponds to the Tyrrhenian 41st parallel line strike-slip fault system (Figure 1a) [Lavecchia, 1988; Spadini and Wezel, 1995]. Both of these fault systems were active during Pliocene Early Pleistocene times, although the Tyrrhenian system is somewhat younger [Bruno et al., 2000]. [8] Significant structural differences have been identified inside the Campania Province. Moho depth increases from the west to the east: the Moho occurs at less than 10 km under the central Tyrrhenian basin [Duschenes et al., 1986; Gueguen et al., 1997]; at approximately 25 km under Ischia Island, between 30 and 35 km below SVV, and at more than 40 km beneath the Southern Apennine 4of25

5 [Chiarabba et al., 2005; Corrado and Rapolla, 1981; Ferrucci et al., 1989]. 3. Geological and Geophysical Data on the CP and Neapolitan Volcanic Region 3.1. Structural and Seismic Data: Fault Patterns and Basement Depth [9] Figure 2 shows the main geophysical data currently available for the CP, together with the fault patterns determined from seismic surveys. A large negative Bouguer anomaly is located within the CP, to the west of the Apennines (Figure 2b). This anomaly can be related to the general deepening of the carbonate basement (which outcrops in the nearby Apennine chain) toward the Bay of Naples and, in general, toward the center of the CP [Bruno et al., 1998, 2003; D Argenio et al., 1973; Finetti and Morelli, 1974; Ippolito et al., 1975]. Four main negative gravity anomalies are centered at Volturno, Acerra, Pompei, and Campi Flegrei caldera. Carbonates have been estimated to occur at more than 3000 m depth in the Volturno basin [Ippolito et al., 1975; Mariani and Prato, 1988; Rosi and Sbrana, 1987], at about 2800 m under the Acerra depression [Bruno et al., 1998], and at 2000 m in the Pompei zone [Bruno et al., 1998]. Only the negative anomaly centered on Pozzuoli may be linked to a depression possibly coinciding with a caldera structure [e.g., Barberi et al., 1991; Florio et al., 1999; Zollo et al., 2003]. Basins situated to the north of the 41st parallel line and to the south of Pompei are shallower (less than 500 m). [10] Analysis of the CP fault pattern allows us to distinguish three main sectors Southern CP [11] In this sector, NE and NW striking normal faults dissect the carbonate basement and possibly can be joined with those cutting the southern Apennines (Figure 2a). For example, the asymmetric graben of Acerra (Figure 2), is related to a NW trending and NE dipping listric fault system. SVV is located at the intersection of NE-SW and NW-SE faults [Bruno et al., 1998]. In particular a NE-trending fault extends offshore of the volcano [Finetti and Morelli, 1974; Bruno et al., 2003], as also suggested by the NE trend of both the Bouguer and magnetic anomaly fields. (Figures 2b and 2c). The influence of regional tectonics on volcanic activity in the PVD is suggested by the alignment of several monogenetic volcanic edifices along NE-SW lineaments (Figures 2a and 5a) [D Antonio et al., 1999a; De Astis et al., 2004; Piochi et al., 1999; Vezzoli, 1988]. NE-SW and NW-SE regional faults also influence the structural evolution of Ischia [Bruno et al., 2002b; Orsi et al., 1991; Sansivero, 1999; Vezzoli, 1988] and the eastern collapsed part of Campi Flegrei [Orsi et al., 1996], as well as the submarine caldera structures in the Bay of Naples [Bruno et al., 2002a] Bay of Naples [12] In this area the faults exhibit a prevailing NE strike, with the exception of the Gulf of Pozzuoli (and PVD), where minor NW and E-W striking faults also occur [Bruno et al., 2003; Bruno, 2004; Finetti and Morelli, 1974; Milia et al., 2000]. Figure 2a shows a main set of NE-SW striking faults and fractures connecting the Magnaghi canyon with the Sebeto fault and defining a NE striking structural discontinuity, named by Bruno et al. [2003] as the Magnaghi-Sebeto line, which cuts through the city of Naples. Many faults/ fractures are located just below the Ischia and Pentapalummo volcanic banks. Important structural differences occur to the south and north of this discontinuity (Figure 3): (1) in the southern sector (SVV), the basement shows an almost regular dip toward NW, with the only exception being the Bocca Grande horst; (2) in the northern sector, the fault arrangement becomes more complex and numerous fractures characterize the PVD offshore [Bruno et al., 2002a, 2003; Bruno, 2004; Finetti and Morelli, 1974] Northern CP [13] This sector is characterized by prevalent WNW and NE striking faults. Inland, the NEtrending fault that delimits the Mt. Massico horst to the south is part of the Ortona-Roccamonfina line (Figure 2a) [Billi et al., 1997; Mantovani et al., 1996; Patacca et al., 1990] and shows features consistent with normal movement [Billi et al., 1997]. Offshore, seismic data by Bruno et al. [2000] highlight the following elements: (1) basins and heights of the carbonate basement match the main structures inland; (2) NE-trending normal faults represent the offshore propagation of the Ortona-Roccamonfina line; (3) WNW to E-W faults on the 41st parallel line are characterized by sinistral strike-slip movements and are the more important structures in the area since they dissect the Ortona Roccamonfina line in the offshore Mt. Massico. These authors describe the 41st parallel 5of25

6 Figure 3. Complex attributes (A, instantaneous amplitude; P, instantaneous phase) of two marine seismic reflection lines, acquired by the Osservatorio Geofisico Sperimentale, Trieste, Italy [Finetti and Morelli, 1974] and reprocessed by Bruno et al. [2003]. Both lines intersect SVV and PVD districts and allow visualization not only of the structural settings of the two areas but also of the greater tectonic complexity of the PVD area with respect to SVV district. line as a deep-seated transfer fault system formed in response to the different rates of opening of various sectors of the Tyrrhenian Sea. Furthermore, strike-slip movements along the 41st parallel faults and normal movements along the Ortona Roccamonfina line are consistent with NW extension [Bruno et al., 2000], which is responsible for longitudinal extension in the Southern Apennine belt [Oldow et al., 1993] Aeromagnetic and Heat Flux Data: Magma Bodies Distribution [14] The aeromagnetic anomaly map (Figure 2c) [AGIP, 1981] (total field) displays good correlations with the main structural features of CP. In particular, magnetic anomalies are higher in the roughly triangular area bounded to the N by the 41st parallel line (including Ventotene and S. Stefano volcanic islands; Figure 1b) and to the SE by the PVD volcanic area. Both the 41st parallel line and the Magnaghi-Sebeto discontinuity are characterized by high magnetic gradients. Magnetic data also show that the Magnaghi-Sebeto discontinuity separates the Bay of Naples into two different sectors: (1) a NW sector, which includes the PVD and displays a series of magnetic anomalies, and (2) a SE sector, where the only important anomaly is located on SVV. Observed anomalies are likely caused by cooled igneous bodies that usually show high susceptibility. Again, the magnetic trends strongly suggest a link between tectonic structures and magma pathways in CP [see also Orsi et al., 1999]. Magmatism appears to be spatially widely distributed in the northernmost offshore sector of CP, whereas the southern sector shows a concentration beneath the SVV. [15] Similar to what is inferred from the magnetic anomalies, heat flow distribution [Della Vedova et al., 1991] is high in the Neapolitan volcanic area and, in particular, there is a NE-trending high heat flow anomaly that perfectly overlaps the PVD alignment (Figure 2d). The elongation of this anomaly is in good agreement with the NE-trending fault pattern characterizing the Gulf of Pozzuoli and the areas offshore of Ischia and Procida. This pattern indicates that 6of25

7 Figure 4. (a) Photo of Somma-Vesuvius stratovolcano during the last 1944 A.D. eruption; (b) geological sketch map of Somma-Vesuvius (modified from Ayuso et al. [1998] and reprinted with permission from Elsevier); and (c) reconstructed volcanological record for the Somma-Vesuvius stratovolcano during the past 25 kyr B.P. Vm, volume of erupted magma (DRE); Vd, volume of deposits. Stratigraphic and volume data from Arrighi et al. [2001]; Cioni et al. [1995]; Civetta and Santacroce [1992]; Cortini and Scandone [1982]; Landi et al. [1999]; Lirer et al. [1973]; Mastrolorenzo et al. [1993]; Rolandi et al. [1993a, 1998]; Rosi and Santacroce [1983]; Santacroce [1987]; Santacroce et al. [1993]; Scandone et al. [1986]. magmatism correlates strongly with the tectonic structures Volcanological Data: Magma Supply System and Output Rate Somma-Vesuvius Volcano [16] SVV is a composite stratovolcano consisting of a recent cone, Vesuvius, which evolved within the older Somma caldera [Santacroce, 1987], built over an area of about 100 km 2 (Figures 4a and 4b). Few parasitic eruptive vents have been recognized on the slope of the volcano [Santacroce, 1987; Bruno and Rapolla, 1999]. The oldest SVV rocks, sampled in boreholes, have been dated at 360 ka [Brocchini et al., 2001], but a detailed record of the SVV activity is limited to the last 25 kyr (Figure 4c). Volcanism consists of quiet effusive eruptions and a spectrum of explosive eruptions. 7of25

8 a Table 1. Geochemical, Thermal, Rheological, and Volcanological Data for Phlegrean Volcanic District and Somma- Vesuvius Phlegrean Volcanic District Somma-Vesuvius CO t/d 150 t/d 3 He/ 4 He 2 5 Ra Ra Thermal gradient 150 C/100 km 15 C/km Temperatures at fumarols and/or thermal C C springs at depth <3 km Viscosity at 4 wt% H 2 O T = 830 C 10 5 Pa s 10 3 Pa s Magma output rate 30 km 3 /10 4 y/10 3 km 3 80 km 3 /10 4 y/10 3 km 3 a CO 2 flux from Chiodini et al. [2001a, 2001b, 2003]. 3 He/ 4 He from Graham et al. [1993]; Tedesco et al. [1990]; Tedesco [1996]; Frondini et al. [2004]; M. Piochi (unpublished data, 1994). Thermal gradient from Rosi and Sbrana [1987], Cataldi et al. [1995], G. Ricciardi (personal communication, 2004), Della Vedova et al. [1991]. Temperatures from Inguaggiato et al. [2000]; Panichi et al. [1992]; Caliro et al. [1999]. Viscosity from Romano et al. [2003]. Magma output rate following Takada [1994a]. Six plinian eruptions have been identified by stratigraphic studies: Codola (25 kyr B.P.) [Santacroce, 1987]; Pomici di Base or Sarno (18 kyr B.P.) [Ayuso et al., 1998; Bertagnini et al., 1998; Landi et al., 1999; Santacroce, 1987]; Novelle/Seggiari-Bosco or Verdoline (16 14 kyr B.P.) [Ayuso et al., 1998; Santacroce, 1987]; Mercato or Ottaviano (8 kyr B.P.) [Rolandi et al., 1993a; Santacroce, 1987]; Avellino (3.5 kyr B.P.) [Cioni et al., 1995; Rolandi et al., 1993b; and references therein]; Pompei (79 A.D.) [Cioni et al., 1995]. Two subplinian eruptions also occurred at 472 A.D. (Pollena) [Rosi and Santacroce, 1983; Santacroce, 1987] and at 1631 A.D. [Rolandi et al., 1993c; Rosi et al., 1993]. Lava effusions and smallscale explosive eruptions took place during periods of nearly continuous volcanism; these interplinian periods [Arrighi et al., 2001; Rolandi et al., 1998; Santacroce, 1987] are the Protohistoric ( ka), the Ancient Historic ( A.D.), the Medieval ( A.D.), and the Recent ( A.D.), when the last SVV eruption occurred. [17] Generally, subplinian and plinian eruptions generated between 0.1 and 4.2 km 3 DRE (Dense Rock Equivalent) of magma [e.g., Cioni et al., 1995; Civetta and Santacroce, 1992; Landi et al., 1999; Rolandi et al., 1993a, 1993b, 1993c; Rosi and Santacroce, 1983; Santacroce, 1987], while 0.01 to 0.4 km 3 DRE of magma were typically erupted during intermediate and small-scale events [Mastrolorenzo et al., 1993; Rolandi et al., 1998; Santacroce et al., 1993; Scandone et al., 1986] (Figure 4c). We have estimated that the total volume of magma erupted in the last 25 kyr is about 20 km 3, corresponding to an average magma output rate of m 3 /yr (Table 1). This value is comparable to that reported by various authors for the A.D. interplinian period ( m 3 /yr) [Cortini and Scandone, 1982; Santacroce et al., 1993; Scandone et al., 1986] and to the magma input rate estimated for the two quiescent periods preceding the eruptions of 472 A.D. ( m 3 /yr) [Rosi and Santacroce, 1983] and of 1631 A.D. ( m 3 /yr) [Rosi et al., 1993]. On this basis, it is possible to assume a roughly constant magma output rate in the last 25 kyr and to infer that the total volume of magma erupted during the lifetime of SVV is probably between 200 and 300 km 3. This is in agreement with data reported by Civetta and Santacroce [1992]. [18] On the basis of the relationship between magma output rate ( m 3 /yr) and event frequency (1 episode each years) during the Recent interplinian period (see Cañon- Tapia and Walker [2004] for a review), and in agreement with the stable position of the vent, we suggest that the SVV system was fed by magma batches rising from depth and percolating through a spatially limited crustal volume. Moreover, on the basis of the morphological features of the volcano and the constant magma output rate we suggest that magma supply system characterizing the SVV throughout its eruptive history remained stable and fairly unchanged; that is, the current plumbing system is similar to that of previous eruptions Phlegraean Volcanic District [19] PVD is an active volcanic area of at least km 2 (Figure 5a) and includes not less than 136 small volcanic edifices such as cinders cones, tuff cones, maars, and domes. The time span of activity ranges from about 300 ka [see De Vivo et al., 2001; Gillot et al., 1982; Rosi and Sbrana, 1987; and references therein] to 1538 A.D. Most of these volcanoes are monogenetic [De Astis et al., 2004; Di Vito et al., 1999; Vezzoli, 1988; and references 8of25

9 Figure 5. (a) DTM (Digital Terrain Model) of the PVD area (from the Laboratorio di Geomatica e Cartografia INGV OV); (b) perspective areal view on eruptive centers at the PVD (from the Laboratorio di Geomatica e Cartografia INGV OV); and (c) reconstructed volcanological record for Campi Flegrei caldera, Procida island, and Ischia island within the PVD during the past 150 kyr B.P. Vm, volume of erupted magma (DRE); Vd, volume of deposits. Stratigraphic, volcanological, and volume data from Acocella et al. [1999]; Beccaluva et al. [1991]; Civetta et al. [1997]; De Astis et al. [2004]; de Vita et al. [1999]; Di Girolamo and Stanzione [1973]; Di Girolamo et al. [1984]; Di Vito et al. [1999]; Finetti and Morelli [1974]; Fisher et al. [1993]; Lirer et al. [2001]; Mastrolorenzo [1994]; Orsi et al. [1992, 1995, 1996]; Pappalardo et al. [1999]; Piochi et al. [1999]; Rosi and Sbrana [1987]; Sansivero [1999]; Scarpati et al. [1993]; Vezzoli [1988]. therein] although a few of them produced multiple eruptions during a limited period time [Di Vito et al., 1999; Isaia et al., 2004; Sansivero, 1999]. Figure 5b shows typical morphology of the PVD, which is due to the presence of multiple craters. [20] The volcanic activity (Figure 5c) mostly consists of effusive and small-scale explosive eruptions, generally extruding less than 0.1 km 3 DRE of magma [de Vita et al., 1999; Di Vito et al., 1987; Lirer et al., 2001; Mastrolorenzo, 1994; Rosi and Sbrana, 1987]. However, the PVD has also experienced very high-magnitude explosive eruptions that deposited huge pyroclastic flow deposits between about 300 and 14 ka [Deino et al., 2004; De Vivo et al., 2001]. Detailed stratigraphic studies have been carried out only on the deposits of the Campanian Ignimbrite (39 ka) [Civetta et al., 1997; De Vivo et al., 2001; Orsi et al., 1996; Rosi et al., 1996; Rosi and Sbrana, 1987; and references therein], and the Neapolitan Yellow Tuff (14 ka) [Deino et al., 2004, and references therein]. The volumes of these deposits are 150 and 20 km 3 DRE, respectively. We agree with previous authors that these well-known pyroclastic sequences originated within the PVD. However, the origin and the emplacement of the Campanian Ignimbrite are still debated [e.g., Di Girolamo, 1970; Orsi et al., 1996; Rosi and Sbrana, 1987; Rosi et al., 1996; Scandone et al., 1991]. In the wake of Di Girolamo [1970] and chrono-stratigraphical data (see De Astis et al. [2004] for a review), we think that the Campanian Ignimbrite was generated at 39 ka by multiple fractures and collapses, in a restricted time span, which drained the same magmatic reservoir. The Monte Epomeo Green Tuff (Ischia), with an age of 55 ka and a volume of 45 km 3 DRE [Vezzoli, 1988] is the third high-energy eruption. The total volume of magma erupted at PVD is 230 km 3 (Table 1) which was generated predominantly during the reported large-scale events. The calculated average magma output rate is m 3 /yr (Table 1). [21] Following Cañon-Tapia and Walker [2004], the event frequency (<1 every 1000 years) and the temporal variation of the position of eruptive vents, combined with the limited dimensions of each volcanic edifice within the PVD, indicate that since the last major eruption the magma supply occurs 9of25

10 Figure 5. (continued) by pervasive intrusions of small magma batches into the crust Compositional and Rheological Data: Magma Viscosity and Accumulation Somma-Vesuvius Volcano [22] The SVV rock crystallinity is strongly variable, from nearly aphyric (mostly in the products from plinian eruptions) to highly porphyritic with lavas having as much as 60% phenocrysts [Trigila and De Benedetti, 1993; Villemant et al., 1993]. Olivine and clinopyroxene dominate the phenocryst assemblage and can be associated with irontitanium oxides, feldspar, leucite, phlogopite, biotite and garnet. Rocks from the last 25 kyr are potassic [Appleton, 1972] and silica undersaturated [Peccerillo, 2003], and range from shoshonite to trachy-phonolite, and from trachy-basalt to tephrite and phonolite (Figure 6a) [Ayuso et al., 1998; Civetta and Santacroce, 1992; Santacroce, 1987]. The high alkali contents of volcanic products strongly influence magma viscosity; for example measured viscosity for phonolitic melt varies from 10 3 and 10 8 Pa s (Figure 6c), depending on water content and temperature [Romano et al., 2003]. Volatiles are dominated by H 2 O and CO 2 that can reach 6 wt% (in the most evolved magmas) and 3500 ppm (in less evolved magmas), respectively [see Cioni, 2000; Marianelli et al., 1999]. Pumice vesicularity is up to 80% with a vesicle number density of 10 6 mm 3 [Gurioli et al., 2005]. [23] Detailed petrochemical studies support the occurrence of recurrent differentiation processes affecting mantle-derived magmas during storage and ascent in the crust: crystal fractionation, phenocryst entrainment, fluid exchange, mixing and crustal contamination have been recognized as the 10 of 25

11 Figure 6. (a) SVV and PVD rock compositions modified from Piochi et al. [2004]. Data sources: Ayuso et al. [1998]; Cioni et al. [1995]; Civetta et al. [1991a, 1991b, 1997]; Civetta and Santacroce [1992]; D Antonio et al. [1999a, 1999b]; De Astis et al. [2004]; De Vivo et al. [1993]; Marianelli et al. [1999]; Orsi et al. [1995]; Pappalardo et al. [1999, 2002]; Piochi et al. [1999]; Santacroce et al. [1993]. (b) Viscosity for Campi Flegrei and Somma- Vesuvius melts from Romano et al. [2003]. most important evolutionary mechanisms [e.g., Ayuso et al., 1998; Civetta and Santacroce, 1992; Civetta et al., 2004; Del Moro et al., 2001; De Vivo et al., 1993; Turi and Taylor, 1976; and references therein]. The variations of radiogenic ( 87 Sr/ 86 Sr = , 143 Nd/ 144 Nd = , 206 Pb/ 204 Pb = ; 207 Pb/ 204 Pb = , 208 Pb/ 204 Pb = ), and stable isotopic ratios (do 18 = %) versus major and trace element contents are in agreement with this interpretation [e.g., Ayuso et al., 1998; Cioni et al., 1995; Civetta and Santacroce, 1992; Cortini and Hermes, 1981]. Several petrological studies together with fluid inclusion data [e.g., Belkin et al., 1985; Belkin and De Vivo, 1993; Lima et al., 2003; Marianelli et al., 1999] demon- 11 of 25

12 Figure 7. Spider diagram of Pearce [1983] for (top) PVD and (bottom) SVV mafic (MgO > 4 wt%) rocks. Data sources: De Astis et al. [2004], Ayuso et al. [1998]. Shaded line indicates the intraplate component of Thorpe et al. [1984]. strated that magma storage occurred at depths of km, 8 10 km and at >12 km beneath SVV. However, at present no geophysical evidence of large (more than 1 km 3 ) magma chambers has been found at depths less than 8 km below SVV [Di Maio et al., 1998; Zollo et al., 1996]. The present magma storage zone has been imaged by seismic tomography at 8 10 km depth [Auger et al., 2001; Zollo et al., 1996, 2002] and is possibly linked to a deeper reservoir that extends down to 30 km [De Natale et al., 2001]. [24] The enrichments of the SVV ultrapotassic mafic rocks (MgO > 4%) in Ba, Sr, Rb, Th, their relative depletion in Nb and Ta, and consequently the rather high LILE/HFSE ratios (Figure 7a), suggest an asthenosphere/lithosphere mantle source, modified by subduction-related fluids [Ayuso et al., 1998; Beccaluva et al., 1991; Peccerillo, 2001, and references therein]. [25] At present, the magmatic system contributes to fumarolic activity observed within the volcanic crater and to diffuse CO 2 degassing around the crater. The emitted gases have a typical hydrothermal composition with H 2 O and CO 2 as major components, followed by H 2,H 2 S, N 2,CH 4 and CO [Chiodini et al., 2001a]. However, the isotopic ratios of helium (R/Ra of 2.7 in Table 1; Ra is the 3 He/ 4 He ratio of the atmosphere ( )) and of carbon (d 13 C from 0 to 0.1%) [Federico et al., 2002] are similar in fumarolic gases and in 12 of 25

13 fluid inclusions from ejecta, indicating a magmatic signature. About 150 tons per day of CO 2 are released from the volcano (Table 1) Phlegraean Volcanic District [26] The PVD rocks contain highly variable amounts (3 60%) of feldspar, clinopyroxene, biotite, apatite, iron oxides and olivine phenocrysts [e.g., D Antonio et al., 1999a; De Astis et al., 2004; Pappalardo et al., 1999; Piochi et al., 1999, 2004; and references therein]. Crystallinity is generally higher in lavas than in pyroclastic rocks and higher in rocks from Ischia with respect to the Campi Flegrei and Procida [Piochi et al., 2004]. [27] Following Appleton [1972], rocks are defined as potassic and range from basalts to trachy-phonolites (Figure 6b) [e.g., Armienti et al., 1983; D Antonio et al., 1999a, 1999b; De Astis et al., 2004; Pappalardo et al., 1999; Piochi et al., 1999, 2004; and references therein]. The lower alkali content of PVD with respect to SVV magmas results in generally higher melt viscosities (Table 1). The viscosity of a representative trachytic melt is between 10 5 and 10 9 Pa s (Figure 6c), depending on water content and temperature [Romano et al., 2003]. The water content is up to 6 9 wt%, and CO 2 reaches 4000 ppm in least evolved magmas [see Signorelli et al., 2001; Marianelli et al., 2002; Cecchetti et al., 2001]. Pumice vesicularity is between 60 and 80% and the vesicle number density ranges from to mm 3 [Mastrolorenzo et al., 2001; Orsi et al., 1992; Piochi et al., 2005]. [28] Major, trace element and Sr-Nd-Pb isotope ratio covariations have been attributed to complex evolutionary processes involving magma chamber refilling, magma mixing and contamination of mantle-derived magmas at various depths in the crust [e.g., D Antonio et al., 1999a; De Astis et al., 2004; Pappalardo et al., 1999, 2002; Piochi et al., 1999]. A large data-set of both radiogenic and stable isotope compositions is available for PVD rocks [e.g., Civetta et al., 1991a, 1991b; D Antonio et al., 1999a, 1999b; De Astis et al., 2004; Pappalardo et al., 1999, 2002; Piochi et al., 1999; and references therein]. Sr isotope compositions range from to for Ischia and Procida, and from to for Campi Flegrei rocks. Nd isotopic ratio ranges between and for Ischia and Procida rocks, whereas the same ratio is between and for Campi Flegrei rocks. 206 Pb/ 204 Pb, 207 Pb/ 204 Pb and 208 Pb/ 204 Pb isotopic ratios vary within the ranges , and , respectively. do 18 varies between 6 and 9% and oxygen isotope ratios are relatively low at Ischia and relatively high at Campi Flegrei [Turi and Taylor, 1976]. [29] Geochemical data indicate that the nearprimary Procida basalts are enriched in LILE, light rare earth elements (LREE), and medium rare earth elements (MREE), and show slight negative HFSE and high rare earth element (HREE) anomalies relative to average Mid-Oceanic Ridge Basalt (MORB) (Figure 7b). (LREE: La, Ce, Pr, Nd, Pm; MREE: Sm, Eu, Gd, Tb, Dy, Ho; HREE: Er, Tm, Yb, Lu.) These features reveal the chemical effects of subduction-related processes, as seen for SVV rocks. However, if Procida basalts are compared with mafic magmas erupted within Campi Flegrei caldera and from SVV, they display lower LILE-REE enrichment and a lower depletion of Ta- Nb. These basalts also have the lowest 87 Sr/ 86 Sr, 206 Pb/ 204 Pb and 208 Pb/ 204 Pb ratios if compared with Campi Flegri and SVV rocks. According to De Astis et al. [2004], Procida basalts can be related to an enriched mantle source with intraplate features, containing subordinate slab-derived components. Thus both trace elements and isotopic data of the mafic rocks (MgO > 4 wt%) indicate that PVD magmatism has been fed by a mantle source variably enriched in incompatible elements and having high Sr and Pb isotope ratios and low Nd isotope compositions [e.g., Beccaluva et al., 1991; D Antonio et al., 1996; D Antonio et al., 1999b; De Astis et al., 2004; Peccerillo, 2001; Piochi et al., 2004]. [30] Mafic rocks from Campi Flegrei caldera were probably stored in a Hercynian- (or Variscan-) type basement and suffered shallow-level evolutionary processes of assimilation, fractionation, and crystallization (AFC) [Pappalardo et al., 2002] which further increased their radiogenic character. Fluid inclusion studies [Cecchetti et al., 2001] confirm this interpretation because they identify differentiation depths >15 km for the Procida mantle-derived magmas and depths between 10 and 15 km for Campi Flegrei trachybasalts. The significant compositional differences associated with chrono-stratigraphical data allowed De Astis et al. [2004] to suggest that different plumbing systems have existed beneath the Ischia-Procida-Torregaveta (southwestern corner of Campi Flegrei, Figure 5) area and the Campi Flegrei caldera since about 40 ka. Magmatic storage also occurs between 3 and 6 km 13 of 25

14 Figure 8. Sketch of magmatic systems and output rate versus deformation rate diagram after Takada [1994a] and considering the models of Gudmundsson [1988, 2002] and Cañon-Tapia and Walker [2004]. Squares indicate polygenetic volcanoes; circles are volcanic fields (data from Takada [1994a]). Also, data for PVD and SVV have been plotted using output rate calculated on the basis of volcanological data (see Table 1), and the deformation rate has been assigned on the basis of geophysical and volcanological evidence. See text for details. depth [De Astis et al., 2004; Marianelli et al., 2002; Piochi et al., 2005; Signorelli et al., 2001]. At the present no evidence for magma bodies underneath the Campi Flegrei with volume larger than 1 km 3 has been found by 3D seismic imaging down to 6 km depth [Zollo et al., 2003; Judenherc and Zollo, 2004]. [31] At present, the PVD is characterized by a high geothermal gradient (up to 180 /km) [AGIP, 1987], widespread fumaroles and thermal springs, as well as by abundant CO 2 degassing which reaches about 1500 tons per day (Table 1). Temperature increases from about 200 C between 500 and 1000 m depth up to C at depth >2000 m beneath Campi Flegrei. High temperatures between 100 and 225 C also occur beneath Ischia in the depth range m [Barbieri et al., 1979]. The chemical composition of emitted gases is dominated by H 2 O and CO 2 [Chiodini et al., 2001a]. However, the isotopic ratio of helium varies between 3 and 2 Ra (see Table 1) [Tedesco et al., 1990; Tedesco, 1996] and reaches its maximum values of about 5 Ra in Forich olivines from Procida (M. Piochi, unpublished 14 of 25

15 data, 1994) [see also Martelli et al., 2004]. Similar ratios in both fumarole gases and fluid inclusions in ejecta, indicate a magmatic signature of the emitted gases. 4. Discussion [32] Volcanological, geochemical, structural and geophysical data suggest that the PVD and SVV plumbing systems are different. A widespread fissure-type system fed by multiple dikes appears to be active beneath the PVD, where NE-SW directed regional faults contribute to the tapping of the least evolved magmas from depths >15 km (i.e., Procida in Figure 5b) [De Astis et al., 2004]. Such a magma supply system (Figure 8, right side) is most consistent with the main geological features of the PVD: (1) a widespread distribution of thermal springs [Caliro et al., 1999; Chiodini et al., 2003; Inguaggiato et al., 2000; Panichi et al., 1992; Panichi and Volpi, 1999] and fumaroles [Chiodini et al., 2001b; Rosi and Sbrana, 1987]; (2) the highly fractured nature of the upper crust as revealed by seismic reflection data (Figure 3) [Bruno et al., 2003]; (3) a generally high heat flow; (4) the highly magnetized nature of the crust as inferred from aeromagnetic data (Figure 2c). All these features suggest pervasive distribution of magma bodies within the crust. [33] By contrast, a long-lived central magma supply system is hypothesized for the SVV (Figure 8, left side). This kind of system is also supported by the existence of a highly rigid [e.g., De Natale et al., 2001; Vilardo et al., 1999] and magnetized [Fedi et al., 1998] crustal volume representing the volcanic conduit that extends from the crater down to about 5 km of depth Monogenetic Versus Polygenetic Volcanism in the CP: Which Mechanism is Dominant? [34] The mechanisms creating fissure-type versus central-type magma plumbing systems are not completely understood. At divergent plate boundaries the style of volcanism depends on the emplacement of magma to form a sill and on the progressive evolution of the system as consequence of the horizontal stress regime above the source region [Gudmundsson, 1988, 1990]. Nevertheless, the regional setting is postcollisional, nested within a dominantly subduction geodynamical regime. Therefore other factors regulating the development of magma supply systems must be considered: magma viscosities, magma supply rate, tectonic regime, initial geometry of fissure systems, and heterogeneities and mechanical proprieties of wall rocks [Cañon-Tapia and Walker, 2004; Fedotov, 1981; Takada, 1989, 1994a, 1994b, 1999; Thorarinsson et al., 1973; Wylie et al., 1999; and references therein]. The most recent study on this topic [Cañon-Tapia and Walker, 2004] provides a general framework for global aspects of volcanism by considering the combined effects of the mantle source, the magma production rate, the regional stress field and the lithology in the upper crustal levels. [35] Field observations [Thorarinsson et al., 1973], theoretical investigations and laboratory experiments (see Wylie et al. [1999] for a review) have shown that localization of magmatic flow can occur as a consequence of the increase of magma viscosity during the ascent through a fissure. However, in the Neapolitan volcanic region, a transition from fissure- to central-type activity cannot be due to a magma viscosity increase (Figure 6). Indeed, Romano et al. [2003] have found a lower melt viscosity for PVD magmas compared to SVV magmas (up to two orders of magnitude; see section 3.4; Figure 6c; Table 1). The higher crystal content of SVV magmas [Trigila and De Benedetti, 1993; Villemant et al., 1993] compared to PVD magmas can lead to an increase in viscosity by about one order of magnitude [Spera, 2000]. However, the anticipated viscosity increase due to crystallinity should be balanced by an opposite effect due to temperature: SVV magmatic temperatures are typically higher (mostly >950 C) than those of PVD magmas (mostly <950 C) [see Belkin and De Vivo, 1993; Civetta et al., 1997; Lima et al., 2003; Marianelli et al., 1999; Signorelli et al., 2001]. In addition, PVD and SVV magmas have similar water and CO 2 content (see section 3.4). Moreover, the relationship between number bubble density and mean equivalent diameter seems to indicate generally lower viscosity for SVV. We conclude that viscosity does not play a major role in magmatic flow localization. [36] Contrary to some authors [Fedotov, 1981; Hildreth, 1981; Walker, 1993], Cañon-Tapia and Walker [2004] suggested that the style of volcanism does not depend on the rate of magma supply. Comparable volumes of erupted magma have in fact been found at polygenetic and monogenetic volcanoes [e.g., Connor and Conway, 2000; Gudmundsson, 1998; Nakamura, 15 of 25

16 1964; Wadge, 1980, 1982]. This fact is also observed in the Neapolitan volcanic region for the entire known period of volcanism (300 ka), and especially during the last 25 kyr (Table 1). For comparative purposes, we normalized the magma output rate to a period of 10 4 years and to a surface area of 10 3 km 2 [Takada, 1994a]. The values we obtained have the same magnitude: km 3 /10 4 yr/10 3 km 2 (SVV) and 30 km 3 / 10 4 yr/10 3 km 2 (PVD). The output rate for PVD magma is also significantly higher than that usually proposed for the generation of monogenetic volcanic fields (less than 1 km 3 /10 4 yr/ 10 3 km 2 ; see examples of Takada [1994a] or Walker [1999]). [37] Cañon-Tapia and Walker [2004], also considering the models developed by Gudmundsson [1988, 1990], suggested that both a high magma pressure within the source, and associated with a magma chamber sited above the source and a constant horizontal minimum stress, can generate monogenetic landforms at the surface. However, this setting, observed at divergent plate margins, is inappropriate for the monogenetic activity in the PVD area. Moreover, these authors [Cañon-Tapia and Walker, 2004; Gudmundsson, 1988, 1990, 1998] suggested that when magma ponding at depth occurs, monogenetic vents can form in association with a central-type volcano. Again, this is not the case for the PVD, as a stratovolcano edifice is missing. Furthermore, PVD and SVV magma chambers formed at similar depth (see section 3.4). [38] On the basis of crack theory, Takada [1994a, and references therein] proposed that the tectonic regime controls the magma supply systems and volcanic styles. Figure 8 shows that polygenetic volcanoes develop in areas characterized by low deformation rates where magmatic pressure is higher than the minimum stress (s1) and allows interaction between liquid-filled cracks (here dykes coalescence) during magma ascent. By contrast, high deformation rates prevent dike coalescence and favor monogenetic centers. This latter condition generally occurs in rift regions where crustal extension balances the stress produced by magma intrusion through high crack propagation rates. [39] The data presented in this paper are largely consistent with a rifting mechanism, and the differences in the style of volcanism between PVD and SVV can be explained in terms of tectonic regime and changing stress conditions between the two areas (Figure 8). Geological and geophysical data [Bruno et al., 2003; D Argenio et al., 1973; Finetti and Del Ben, 1986; Ippolito et al., 1975] indicate the existence of an extensional tectonic regime at the PVD. Airborne gravity data [Rosi and Sbrana, 1987] show that the deeper basins are located in the central and northern sectors of CP (Figures 2a and 2b), which are also characterized by high values of magnetic anomalies. Seismic data confirm that the top of the carbonate basement occurs at higher depths in these sectors of the CP [Bruno et al., 2000, 2002a, 2003; Finetti and Morelli, 1974; Zollo et al., 2003]. The increase in crustal thickness [e.g., Corrado and Rapolla, 1981; Ferrucci et al., 1989] from the PVD (around 20 km) to the SVV (around 30 km) also supports the different eruptive behavior of the two areas. Moreover, in the PVD, these extensional processes are mirrored in the alignment of several monogenetic volcanic edifices along NE-SW and NW-SE regional faults [Bruno et al., 2003; Orsi et al., 1991, 1996; D Antonio et al., 1999a; De Astis et al., 2004; Piochi et al., 1999; Sansivero, 1999; Vezzoli, 1988]. Furthermore, there are some compositional features of the Neapolitan mafic rocks that are consistent with our interpretation. The trace element geochemistry of these rocks supports the idea of a within-plate component in the genesis of the most primitive magmas erupted at Procida [e.g., De Astis et al., 2004; Piochi et al., 2004, and references therein], and indicates distinct mantle sources for PVD and SVV volcanic rocks, as we discuss later on. [40] In summary, we propose that an extensional tectonic regime is active within the PVD, while extension at SVV is weak. This polygenetic volcano is located just outside the high deformation zone affecting the PVD and on a less dissected and shallower carbonate basement. These distinct geological situations can explain the development of different magma supply systems at SVV and PVD which are only separated by a distance of 10 km (Figure 1b) Tectonic Setting and Kinematics of Campanian Plain [41] Structural, seismic, Bouguer and aeromagnetic anomaly data for the CP (Figures 2a, 2b, and 2c), allow us to define the boundaries of the PVD extensional area and to differentiate it from the SVV area. The roughly EW-striking 41st parallel 16 of 25