Understanding Magma Evolution at Campi Flegrei (Campania, Italy) Volcanic

Transcription

1 1 2 Understanding Magma Evolution at Campi Flegrei (Campania, Italy) Volcanic Complex Using Melt Inclusions and Phase Equilibria 3 4 Cannatelli C. a, *, Spera F.J. a, Fedele L. b, De Vivo B. c a Department of Earth Science and Institute for Crustal Studies, University of California, Santa Barbara, CA USA b Department of Geosciences Virginia Tech, 4044 Derring Hall, Blacksburg, VA USA c Dipartimento di Scienze della Terra, Università di Napoli Federico II, Napoli, Italy * Corresponding author: Tel , Fax: addresses: claudia@crustal.ucsb.edu (C. Cannatelli), spera@geol.ucsb.edu (F.J. Spera), lfedele@vt.edu (L. Fedele), bdevivo@unina.it (B. De Vivo)

2 Abstract The magmatic evolution of two eruptive episodes at Campi Flegrei (Italy) has been investigated using phase equilibria modeling (MELTS) and data from melt inclusions (MIs) in phenocrysts from the Fondo Riccio (FR, 9.5 ka) and Minopoli 1 (Mi1, 11.1 ka) eruptions. Adopting the Ansatz that isobaric fractional crystallization of a mantle-derived parental magma is the dominant petrogenetic process, major element evolution and corresponding changes in the physical and thermodynamic properties of the magma bodies from which FR and Mi1 magmas were erupted can be tracked. Using olivine hosted MIs as representative of parental melt, the physical conditions and crystallization path have been modeled. Results are compared to observed crystal, whole rock and homogenized MI compositions to evaluate the extent computed phase equilibria can reproduce observations under the imposed conditions. FR parental magma was likely trachyandesitic, approximated by the composition of MIs in olivine (SiO 2 = 46.8%, MgO = 9.45 %), which evolved mainly through fractional crystallization at low pressure (P 0.2 GPa, 8 km depth), along the QFM±1 oxygen buffer with an initial dissolved H 2 O content of ~3 wt%. Mi1 parental magma was also trachyandesitic and it is approximated by the chemistry of MIs in olivine (SiO 2 = 47.8%, MgO = 9.37%). The estimated mean pressure of crystallization is 0.3 GPa ( 12 km depth), deeper than FR with oxygen fugacity along QFM+1buffer. The initial H 2 O content of ~ 2 wt% for Mi1 is slightly less than that of FR. Thermodynamic modeling also suggests that mafic parental magma crystallized by about 50% to generate the more evolved (erupted) compositions. MIs in olivine phenocrysts, the first phenocryst to crystallize, evidently represent trapped pristine remnants of the parental magma. MIs within later formed clinopyroxene phenocrysts do not appear to represent equilibrium liquids trapped along the liquid line of descent suggesting that reaction between trapped melt and clinopyroxene may be important or that significant liquid heterogeneity developed by the time clinopyroxene began to crystallize. The relationship between melt fraction and T reveals for FR the presence of a pseudo-invariant temperature, T inv = 880 at which the fraction of melt decreases abruptly due to simultaneous crystallization of alkali feldspar and plagioclase, eutectic- 2

3 like behavior. The melt density, viscosity and dissolved water content change abruptly in a very small temperature interval around T inv. At this temperature, the volume fraction of exsolved H 2 O present within magma increases from less than 10% by volume to more than 60 vol % which is of the order of the fragmentation limit of circa 60 vol% for FR differentiated parent melt. In the case of Mi1, simulations do not point to abrupt invariant temperature behavior but instead melt fraction (f m ) varies from 0.5 to 0.2 in a temperature span of 90 C (around 990 C), due to the crystallization of alkali feldspars, plagioclase and biotite. This less eutectic-like behavior may be due to higher mean crystallization pressure of Mi1 compared to FR. A simple thermal model based on variation of enthalpy of the system along the liquid line of descent allowed us to estimate the duration of the entire differentiation event, suggesting a timescale for FR of 6.5 ± 3.5 kyr and for Mi1 of 2.5±1.5 kyr from the beginning of fractionation until eruption. 62 3

4 Introduction Campi Flegrei (CF) (Italy) is the most active magmatic system in the Mediterranean region and has exhibited predominantly explosive volcanic activity for more than 300,000 years (Pappalardo et al., 2002). The area is well known for its intense hydrothermal activity, frequent earthquakes and long history of bradyseism including the recent episodes in and The city of Naples and surroundings, with ~4 million inhabitants, represents one of the most densely populated and volcanically active areas on Earth. The origins of CF s explosive volcanism have been the focus of intense research for hundreds of years and is still debated today (Di Girolamo et al, 1984; Rosi and Sbrana, 1987; Barberi et al., 1991; Pappalardo et al., 1999; De Vivo et al., 2001; Rolandi et al., 2003; De Astis et al., 2004; De Vivo and Lima, 2006; Marianelli et al. 2006; Bodnar at el., 2007; Di Vito et al., 2008; Lima et al., 2009). Explosive volcanic eruptions constitute a challenge for volcanologists because of their unpredictability; identification of the parameters determining the style of an eruption is of fundamental importance in efforts to understand how explosive volcanoes work. Development of models for volcanic eruption forecasting require information on the pre-eruptive chemical and physical characteristics of the magmatic system (Anderson et al., 2000; Webster et al., 2001; Roggensack et al., 2001; De Vivo et al., 2005; Metrich and Wallace 2008; Moore 2008). In particular the pre-eruptive composition of the magma before the eruption, including its dissolved volatile content, is of critical importance because composition exerts a fundamental control of magma properties and hence the style of eruptive events (Anderson, 1976; Burnham, 1979; De Vivo et al., 2005). The exsolution and expansion of volatiles (especially H 2 O) provides the mechanical energy that drives explosive volcanic eruptions. The physical properties of magmas, such as density and viscosity, (Lange 1994; Ochs and Lange, 1999; Spera et al, 2000) along with the pre-eruptive phase equilibria (Moore and Carmichael, 1998) are strongly influenced by the dissolution of volatiles in magma and affect the volcanic style of a magmatic system (Sparks et al., 1994). 4

5 Melt inclusions (MIs) are a powerful tool to investigate the pre-eruptive magma composition since they potentially retain the pristine composition of the magma at the time of trapping (Roedder 1984). The original volatile content of magma can be estimated by analyzing MIs contained in phenocrysts (Anderson, 1974; Clocchiatti, 1975; Roedder, 1979; Belkin et al., 1985; Sobolev, 1990; Lowenstern, 1994; Anderson, 2003; De Vivo and Bodnar, 2003; Wallace, 2005). Moreover, MIs provide information concerning crystallization and mixing histories of magmas and also the conditions of primary melt generation and extraction (Roedder, 1984; Carroll and Holloway, 1994; Lowenstern, 1994; Sobolev, 1996; Danyushevsky et al., 2000; Frezzotti, 2001). In the past two decades great of effort have been devoted to the description of the processes that drive the evolution of sub-surface magmas at Campi Flegrei as well as the eruptions themselves. In particular, some authors (Civetta et al., 1997; Pappalardo et al. 2002; Tonarini et al. 2004; Roach, 2005; D Antonio et al., 2007; Arienzo et al., 2009; Tonarini et al., 2009) have shown that fractional crystallization, magma mixing and perhaps wallrock assimilation also play roles in describing the evolution of CF. Indeed sorting out which of these and possibly other mechanisms is most important is a significant part of petrologic research on the evolution of crustal magma bodies. In the present work we examine the origin of magma erupted during the Fondo Riccio, FR (9.5 ka) and Minopoli 1, Mi1 (10.3 ka) volcanic episodes by deriving constraints imposed from phase equilibria embodied in the MELTS thermodynamic model (Ghiorso and Sack, 1995), from phenocryst and glass compositions and from an analysis of MIs found in phenocrysts. Using olivine hosted MIs as representative of parental melt that generated the eruptive products of FR and Mi1, estimates of the pressure, temperature, oxygen buffer, density and viscosity can be made assuming isobaric fractional crystallization was the dominant process of geochemical evolution. Although it would be easy to perform polybaric crystal fractionation (indeed other paths could be chosen) the approach here is to adopt the very simplest case and compare detailed predictions to observations. The deviations from the model and observation then put some constraints on the importance of other processes of petrologic evolution. An important aspect of our findings is the identification of a 5

6 pseudo-invariant temperature (T inv ) along the liquid line of descent. At this temperature the system undergoes dramatic changes in crystallinity, melt composition including volatile content, viscosity, and density. The net effect of these changes is to drive the system towards dynamic instability, which we speculate is the trigger mechanism for the eruptions. A simple thermal model based on the variation of enthalpy of the system along the liquid line of descent is also presented to estimate the timescale between the start of significant crystallization and the time of eruption Volcanological background Campi Flegrei Volcanic District (CFVD) is a large volcanic complex (~ 200 km 2 ) located west of the city of Naples, Italy (Fig.1). Multiple eruptions have occurred in this area in the last 300 ka (Pappalardo et al., 2002), as well as intense hydrothermal activity, bradyseismic events and frequent earthquakes. The major eruption occurring in the CFVD is the 15 ka Neapolitan Yellow Tuff (NYT) (Deino et al., 2004). The origin of the Campanian Ignimbrite (CI) (39 ka) is controversial: for some authors (Rosi and Sbrana, 1987; Orsi et al., 1996) this eruption occurred in the CFVD; other authors (De Vivo et al., 2001; Rolandi et al., 2003) suggest that the CI originated from fractures activated along the neotectonic Apennine fault system parallel to the Tyrrhenian coastline. They argue that eruptions from >300 ka to 19 ka are not confined to a unique volcanic center or isolated vent system in CF as suggested by Rosi and Sbrana, 1987 and Orsi et al., De Vivo et al (2001) and Rolandi et al., (2003) argued that only the Neapolitan Yellow Tuff (NYT) (15 ka, Deino et al., 2004) erupted from vents within CF, whereas the CI (39 ka, DeVivo et al., 2001) has a much wider source and dispersal area. According to Pappalardo et al. (2002), the interval between the CI and NYT eruptions is characterized by a number of small magnitude volcanic events. Since the NYT eruption, the margins of the region have been the site of at least 65 eruptions, divided in three periods of activity. Eruptions were separated by quiescent periods marked by two widespread paleosols (Di Vito et al., 6

7 ). The last eruption in 1538 A.D. formed the Monte Nuovo cone (Di Vito et al., 1987) after 3.4 ka of dormancy. In this paper we analyze the Fondo Riccio (FR) and Minopoli 1 (Mi1) eruptive products in an effort to deduce their petrogenesis. The FR eruption was explosive with a strombolian character and occurred at kyr (D Antonio et al., 1999) from an eruptive centre on the western side of the Gauro volcano, near the centre of the Phlegrean caldera (Fig 1). The eruptive deposits are limited to the vent area and lie above the Paleosol A and below the Montagna Spaccata Tephra. The eruptive products consist of fallout deposits composed of very coarse scoria beds with subordinate coarse ash beds (Di Vito et al., 1999). According to Di Vito et al. (1999), the earlier Mi1 eruption occurred ka and was strombolian with subordinate phreatomagmatic phases, while Di Girolamo et al. (1984), based on the degree of dispersal of Mi1 s products, define this eruption as sub-plinian. The deposits are limited to the vent area formed by scoriae horizons with a composition varying from latitic to alkalitrachytic. The eruptive products are composed of alternating pumice lapilli fallout and mainly massive ash fallout beds and, subordinately, cross laminated ash surge beds, rich in accretionary lapilli (Di Vito et al., 1999). Evidently, the Mi1 eruption had a stronger phreatomagmatic component than the closely related FR eruption based on observed stratigraphy Sample description and analytical technique The locations of the samples utilized in this study are indicated in Figure 1. Here we give petrographic and mineralogical descriptions of the samples and describe the methods used to perform the analyses Petrography and chemical composition of Fondo Riccio 7

8 For FR, CF-FR-C1 was collected at the top of the stratigraphic column and is a wellvesciculated scoriae containing approximately 20% by volume of phenocrysts. The phenocrysts include olivine, clinopyroxene, spinel (magnetite), biotite, alkali feldspar and plagioclase. Biotite occurs as large crystals (typical size ~ 2-3 mm), while apatite phenocrysts occur as small (~ 0.1 mm) acicular needles. Clinopyroxene and feldspar commonly exhibit intergrowth textures, suggesting cotectic crystallization. Olivine, clinopyroxene and plagioclase contain recrystallized MIs, while alkali feldspar phenocrysts contain apatite inclusions. Sample, CF-FR-C2, is a bomb, relatively unvesciculated, containing olivine, clinopyroxene, apatite, spinel, biotite, alkali feldspar and plagioclase. Olivine, clinopyroxene and alkali feldspar phenocrysts contain recrystallized MIs. Petrochemically, both samples are porphyritic latite with ~ 20% phenocrysts, with clinopyroxene and plagioclase often found in glomeroporphyritic clots; clinopyroxene and plagioclase also occur as microlites in the groundmass. In the FR samples, olivine phenocrysts range Fo 84-87, and pyroxene lies in the diopside-salite field on the pyroxene quadrilateral, with Wo and Fs 6-15 (Table 1). Based on microprobe analyses, alkali feldspars in FR present a unimodal distribution with Or component of ~ 79 to 88. Plagioclase crystals are zoned with An (Table 2) Petrography and chemical composition of Minopoli 1 For Mi1, CF-MI1-C1 was collected in the Casalesio area (Fig 1), at the base of the deposit. The sample is greyish-black scoriae, of trachybasalt composition containing ~ 20% phenocrysts of olivine, clinopyroxene, plagioclase, alkali feldspar, spinel (magnetite), apatite and biotite. Olivine phenocrysts are weakly to unzoned with average Fo content ~ 78, while pyroxenes present Wo values between 45 and 48 and Fs between 6 and 16 (Table 3). Based on microprobe analyses, alkali feldspars in Minopoli 1 present a unimodal distribution of Or values which ranges from ~ 75 to 80. Alkali feldspars exhibit zonation, with potassic cores. Plagioclase crystals are highly zoned 8

9 presenting a bimodal distribution with a range from ~ 54 to 87 with peaks at 54 and 83 based on about 50 grains (Table 4) Melt Inclusions description The MI s present in both FR and M1 are generally devitrified and partially recrystallized, present a bubble (shrinkage ± exsolution of volatiles) and daughter minerals (generally apatite and oxides). MIs generally have elongated ellipsoidal shapes and range from 30 to 80 µm (most between 20 and 50 µm). In order to be analyzed, MIs needed to be re-heated to a homogenous glass. Detailed descriptions of MIs reheating procedures, sample preparation and analytical methods are in Cannatelli et al., Analytical methods Major and minor elements analyses of phenocrysts were performed in the Department of Earth Science at UCSB using a Cameca SX-50 electron microprobe equipped with five wavelength dispersive spectrometers. Phenocrysts analyses were performed using a 1µm focused beam at 15 kev accelerating voltage and a beam current of 15nA. Uncertainty of analyses is around 1% (relative) for most elements. Quantitative electron microprobe analyses (EMPA) on phenocrysts and MIs were performed at Virginia Tech and at University of Rome La Sapienza (IGAG-CNR, Rome, Italy) on a Cameca SX-50 equipped with four wavelength dispersive spectrometers. The analytical scheme chosen for MIs is described in Cannatelli et al., 2007 and reference therein Phase equilibria modeling Procedures to select the parental melt composition Phase equilibria modeling has been carried out using the software MELTS, a thermodynamic model of crystal-liquid equilibria. The MELTS algorithm is based on classical equilibrium 9

10 thermodynamics and has been object of extensive reviews (Ghiorso and Sack, 1995, Asimow and Ghiorso, 1998). The use of MELTS to reconstruct the crystallization path of a magma requires specification of initial conditions, including 1) the initial state of the system (parental melt composition including H 2 O content, starting temperature and pressure, and oxygen fugacity) and 2) constraints under which the magmatic evolution proceeds (open or closed system, fractional or equilibrium crystallization, minimization of appropriate thermodynamic potential based on imposed constraints). In this work we investigate isobaric crystallization scenarios and explore both equilibrium and fractional crystallization scenarios. The search of parental melt composition starts with the assumption that MIs within phenocryst phases can be related to a unique parental melt during cotectic (olivine +clinopyroxene) crystallization. The graphical method developed by Watson (1976) is used to test the hypothesis that MIs are primary or nearly so. MIs composition(s) of interest are further culled by selecting ones that exhibit the lowest concentrations of incompatible trace elements and highest MgO contents as input for the phase equilibria calculations. In the case of Fondo Riccio, 7 MIs were selected, hosted in olivine and pyroxene and have been plotted on a CaO-MgO-Al 2 O 3 coordinates, as described by Watson (1976). The intersection I (Figure 2a) of olivine and clinopyroxene fractionation lines is in the field occupied by FR-C1-o6 M1, a MIs hosted in olivine O6. This MI represents the predicted composition of the melt at the cotectic point, where olivine and clinopyroxene crystallize simultaneously, so it is reasonable to hypothesize that the Parental Melt (PM) composition should be more primitive than FR-C1-o6 M1. The MIs FR-C1-o2 M1 (9.45 wt% MgO), and FR-C1-o1 M1 (8.05 wt % MgO) possess high MgO contents and the lowest concentration of incompatible trace elements and are consequently considered the best candidates to represent the parental melt. We carried out phase equilibria calculations using FR-C1-o1 M1 (not shown) and FR-C1-o2 M1 and differences were small; based on this we decided to select the one with the highest MgO content. In the case of Mi1, by applying the Watson graphical method we found that Mi1-C1-P8 M1, a MI hosted in the clinopyroxene P8 (fig 2b) represents the composition of the melt at the cotectic point. 10

11 We selected the parental melt composition choosing the MI with the highest MgO content and lowest incompatible trace element concentrations as an approximation to the PM. The MI that best fit the criteria and was closest to Mi1-C1-P8 M1 in Fig. 2b was hosted in olivine o5 with a MgO content of 9.37 wt%, and values of Ce, and Nd of 69 and 61ppm. It is probable that MIs in olivine can undergo some re-equilibration with the host (Danyushevsky et al., 2000; Kress and Ghiorso, 2004). However in our case the MELTS results agree very well with the compositions for the MIs in olivine for both FR and Mi1 samples. Our interpretation of these relations is that that post entrapment changes for these MIs are small to negligible. We conclude that the method espoused 35 years ago by Watson is indeed useful and that by careful use of MIs one can at least in this case estimate the parental melt composition reasonably well Phase equilibria: constraints and limitations To reconstruct the magmatic evolution the initial state of the system, devolatilized PM composition, dissolved H 2 O content of PM, initial temperature, pressure, and oxygen buffer are specified. Here we present results of closed system isobaric fractional crystallization where the Gibbs energy is the appropriate thermodynamic potential to be minimized. We have adopted these constraints as an Ansatz to be tested by the closeness of the computed results to observations. These runs clearly show the effects of varying pressure, f O2 and the initial water content of the parental melt on the liquid line of descent and on the composition and abundance of all crystalline phases and the temperature at which melt becomes water saturated. After setting P, f O2 and dissolved H 2 O content, we compare predicted phase and melt compositions to those observed in order to determine the range of physical conditions leading up to eruption for FR and Mi1. We selected the best case based on correspondence between mineralogical and geochemical data and the phase equilibria calculations. Calculations were rejected when the deviation between observation and model was deemed too large. Although the degree of closeness could be quantified by, for example, using a Euclidean norm criterion comparing the predicted oxide composition of a phase to its observed 11

12 values, we believe such a procedure is a premature at the present time. Instead, we prefer to rely on reasonable judgment predicated on the assumption that an experienced petrologist will be able to spot a poor solution as one that provides no new insight into the petrogenesis of the system and on what may have triggered the eruptions. One must keep in mind the assumptions of the method and the realities of Nature. For example, the calculation assumes perfect fractional crystallization. However, in situations where crystals are removed from liquid by some physical process driven by gravity (e.g., crystal settling/floatation) or deviatoric stress (e.g., kneading, melt percolation, filter pressing, see Kohlstedt and Holtzman, 2009), there will always be some reaction between earlier formed crystals and ambient liquid. Similarly, the calculation assumes there is a single parental composition from which all differentiated liquids develop. It is easy to imagine that compositional heterogeneities would be present a priori even if convective mixing was reasonably efficient. Finally, the calculation assumes that crystallization is isobaric, exactly. The approximate nature of this assumption should be clear to anyone who ever mapped plutonic in rugged terrain. The point of performing phase equilibria calculations using an imperfect thermodynamic model (no thermodynamic model is perfect) with constraints that are clearly approximate is to evaluate the overall reasonability of the proposed scenario. If, for example, crystallization is grossly polybaric, then no isobaric model will come close to reproducing observed phase compositions, abundances and glass (melt) compositions. One could then perform a constrained polybaric simulation and ask if that procedure produces better agreement. If assimilation plays an important part of the petrogenesis, then no closed system phase equilibria model will produce satisfactory correspondence to observations and one would seek to explore alternative petrogenetic models quantitatively involving assimilation and the mixing of melts or magmas of differing composition and temperature. In this study we find that isobaric closed system fractional crystallization at low pressure produces results that bear a close (but not perfect) correspondence to observed relations and that the 12

13 implications of the calculation suggest a causative link between crystallization and the eruptive episode that generated the two small volcanic deposits of the FR and Mi1 (see below) Fondo Riccio The initial water content in the parental melt has been estimated starting from the values obtained for MIs by SIMS analyses. FR s MIs belong to two different populations of inclusions, one with water contents ranging between 1 and 4 wt% and the other with water values around 6 wt%. As starting water content we tested values ranging between 1 and 5 wt%, but from petrographic observations values of H 2 O >3wt% were discarded because of the high water saturation temperature. For example, in the case of 4wt% H 2 O the temperature of water saturation was 1070 C at 0.2 GPa (depth ~ 6 km). At this temperature the system is saturated in water and crystallizing mineral phases such as clinopyroxene, plagioclase and alkali feldspar should trap fluid inclusions during the cooling process. There is no petrographic evidence of fluid inclusions hosted in these phases in the samples studied here. In the cases of H 2 O < 2 wt%, each run generated a rhombohedral oxide phase (illmenite) at low melt fractions, inconsistent with the phase assemblage observed. Although not shown, calculated runs with initial water content in the PM less than 2 wt% and greater than 4 wt% did not predict the phase assemblage observed in the FR. We therefore conclude that initial water content in the PM around 3 wt% is the most realistic case for the FR eruptive system. Although we acknowledge that this is a judgment, we believe it to be the best estimate based on the congruence between calculation and what is observed in the natural samples studied in the laboratory. The majority of the runs were made isobarically and for FR at P < 0.3 GPa; at greater P the presence of predicted minerals such as garnet or muscovite is not compatible with the FR phenocryst assemblage. To understand better the effect of changing pressure, we compared MELTS generated TAS diagrams calculated at a fixed f O2 = QFM+1, QFM and P = 0.1, 0.15, 0.2 and 0.3 GPa. For the case of f O2 = QFM and QFM+1 good agreement between phase equilibria (MELTS) 13

14 with FR s data (see Fig. 3). The best case scenario of oxygen fugacity for FR was chosen for P 0.2 GPa, corresponding to ~8 km depth, and compatible with recent studies by Zollo et al., 2008 suggesting that a hypothetical magma body at Campi Flegrei is at least 7.5 km deep. From petrographic investigation we found the presence of spinel (in the form of magnetite solid solution) in olivine and clinopyroxene, but not in plagioclase and feldspars. Biotite is also present. We compared several MELTS generated mineral distribution diagrams with petrographic observations and found best agreement is reached for f O2 between QFM-1 and QFM+1. We also noticed, as expected, the strong dependence of the iron-bearing phases on the variation of oxygen fugacity. For example, when we consider the case of FR with initial water content of 2 wt%, an increase in the oxygen fugacity from QFM-2 to QFM+2, stabilizes spinel at higher temperature, while not affecting the crystallization temperature of clinopyroxenes and feldspars (Fig. 4). The stabilization of spinel at higher temperatures corresponds to a decrease of FeO tot and increase of SiO 2 content in the melt. The inconsistency between observed mineral assemblage and MELTS generated mineral distribution has lead us to discard oxygen fugacity extreme values of QFM-2, QFM-1 and QFM+2. In summary, the physical conditions that produce the closest correspondence between the model and observation is fractional crystallization of a parental melt of (anhydrous) composition (given in Table 5) plus 3 wt % H 2 O added at 0.15 GPa and oxygen fugacity around the QFM buffer Minopoli 1 Water contents of MIs from the Mi1 eruptive products were measured by SIMS and range from 1 to 4 wt% (Cannatelli et al., 2007). The effect of varying the initial water concentration in the parental melt was examined in the Mi1 case through isobaric fractional crystallization as for FR. Petrographic studies of Mi1 s thin sections reveal the presence of large (1-2 mm) biotite crystals. The presence of such crystals implies initial water contents greater than 2 wt%. Therefore simulations obtained by setting the water content less than 2wt% were discarded, regardless of 14

15 oxygen fugacity and pressure values. Furthermore, in the case of H 2 O > 2wt% we observed a lack of intersection between the MELTS generated oxides trends and the real data field of Mi1. In particular, values of water content greater of 3 wt% were discarded for f O2 = QFM QFM+2 and pressure greater than 0.3 GPa, because of the predicted presence of garnet and leucite, inconsistent with the observed assemblage. Values of water greater than 4 wt% were discarded because of the high water saturation temperature (T ~ 1080 C) which would result in the presence of fluid inclusions in the phenocrysts of Mi1 sample, not observed in Mi1. The initial water content of the parental melt for Mi1 estimated is therefore around 2 wt% a bit lower than that FR using the same methods. Several simulations were carried out using a fixed value of initial water content of 2-3wt%, and varying the pressure and the oxygen fugacity. Many runs were discarded because of mismatch between observed and predicted phases, such in the cases of f O2 > QFM or P 0.1GPa. A small decrease in oxygen fugacity leads to a decrease of spinel stabilization temperature of almost 100 C and a longer crystallization interval for feldspars with a consequent greater generated mass of feldspars in the mineral assemblage. Comparisons among feldspars plotting model results and observations on ternary diagrams (An-Ab-Or) and spinel diagrams (FeO-Fe 2 O 3 -TiO 2 ) were studied in order to establish the best fit. In general higher pressures better match observed phases (Fig. 5). In particular, for f O2 = QFM+1 we obtain a good fit for spinel and feldspar compositions to observed M1 phenocryst (see Table 6), especially at P=0.3GPa and water content of 2 wt%. The best case chosen from the several M1 simulations is represented by a parental melt of (anhydrous) composition (Table 7) at pressure P ~ 0.3 GPa (~12 km depth), water content of 2 wt % along the QFM+1 oxygen buffer Results Fondo Riccio 15

16 We present results for FR for isobaric fractional crystallization of the estimated parental composition. In fact, we have used a number of possible parental compositions and although small differences in results are obtained, the salient features are not significantly affected. The parental melt composition of FR-C1-o2 M1 with an initial water content of 3 wt% is used to generate the results below. The fractional crystallization path along the QFM to QFM+1 oxygen buffer at 0.15 GPa has been computed. MELTS correctly predicts the mineral phases observed. Olivine is the liquidus (T= 1260 C) phase, followed by clinopyroxene, magnetite, H 2 O, plagioclase, Alkali feldspar and biotite at 1110 C, 1100 C, 1070 C and 880 C respectively. Mineral distribution, abundances and temperature at which water saturates are shown in Fig. 3. It is interesting to note the abrupt change in melt composition around 880 C due to a simultaneous crystallization of alkali feldspar, plagioclase and biotite. We can define this temperature as a pseudo-invariant point temperature (Fowler et al., 2007). At this fixed temperature, a major change in melt fraction (f m ), from 0.5 to 0.1 and melt composition occurs (Fig. 6) as crystals precipitate from the melt due to extraction of heat (enthalpy). The properties of the melt (density and shear viscosity) and of magma (density, volume fraction of bubbles, shear viscosity) change dramatically around this temperature (see below). The growth of alkali feldspars and plagioclase dominates the crystallization path at T ~ 880 C and below. In fig. 4 crystallization patterns for f O2 = QFM and QFM+1 are portrayed. Concentrations of SiO 2, K 2 O, Na 2 O and Al 2 O 3 initially increase with decreasing MgO due to the crystallization of olivine and continue to increase as clinopyroxene, spinel and apatite crystallize (Fig. 7a-f). The increase of CaO concentration ends when melt becomes saturated in clinopyroxene and then decreases slowly with cooling. FeO tot concentrations slightly decrease in the early stages of crystallization and decrease abruptly when spinel joins the already fractionated phase minerals. Results for QFM and QFM+1 are quite similar for Al 2 O 3, K 2 O and Na 2 O while for SiO 2, FeO tot and CaO we can observe a more close approximation to the observed trends for f O2 =QFM+1. 16

17 At T=T inv there is a rapid change in the variation diagram trajectories of circa 2 wt% for SiO 2 and Al 2 O 3, 1 wt% for K 2 O and 0.5 wt% for CaO and Na 2 O. For T<T inv SiO 2, CaO, Al 2 O 3 and K 2 O show a sudden decrease, while Na 2 O continues to increase as a result of feldspar fractionation. These compositional changes at T=T inv are associated to a change in the physical properties of both melt and magma with significant consequences for eruption probability and dynamics. As noted on Fig. 7, MI s hosted in olivine and clinopyroxene agree well with the predicted liquid line of descent making melt inclusions, especially hosted in olivine. There is a good agreement between observed and simulated clinopyroxene and olivine compositions (Fig. 8) remembering that calculated values assume perfect fractional crystallization. Alkali feldspar trends compare favourably; predicted plagioclase becomes more sodic than observed values near the solidus presumably related to the breakdown of the assumption of perfect fractional crystallization near the solidus where the separation of melt from surrounding solids is slowed, melt percolation through a mush being much slower than crystal settling in a crystal-poor magma. We tested two cases of fractionation: (1) All crystals and exsolved H 2 O (water bubbles) are removed immediately upon saturation and (2) only precipitated crystals are fractionated; water bubbles remain suspended in the melt. The best agreement was found when both precipitated solids and exsolved H 2 O were removed in perfect fractional crystallization. The spinel ternary diagram also shows good agreement between MELTS predictions and spinel compositions from FR samples (Table 5) Minopoli 1 Based on the assumption that the most realistic parental melt has a composition of Mi1-C1-o5- with water content of 2-3 wt%, we present results for calculations with oxygen buffer set at QFM+1 and pressure at GPa. In Fig. 9 we can see the mineral distribution along the crystallization path for the case P= 0.3 GPa, 2wt% H 2 O and f O2 = QFM+1. 17

18 The liquidus phase is olivine at T= 1300 C, followed by clinopyroxene (T= 1160 C), spinel (T=1070 C) and apatite (T=1020 C). At 990 C, 960 C and 900 C respectively plagioclase, biotite and alkali feldspar join the mineral assemblage, dominating the crystallization path. From Fig. 10a-f, concentrations of SiO 2, Al 2 O 3, K 2 O and Na 2 O increase with decreasing MgO during the crystallization of olivine and then continue to increase as clinopyroxene, apatite and spinel crystallize. The increase of CaO ends when clinopyroxene begins crystallization and then decreases slowly with cooling. FeO tot slightly decreases in the early stages of crystallization, then remains constant and decreases abruptly only when spinel begins fractionation. In the case of Mi1 an abrupt change in melt composition noted in FR is not evident; instead, in a temperature span of about 80 C (around T= 990 C) there is a change of f m from 0.5 to 0.2, due mainly to the crystallization of feldspar. At T=990 C, there are changes in the calculated oxides trends of about 3 wt% for SiO 2, CaO and K 2 O, 2wt% for Al 2 O 3 and 1 wt% for Na 2 O and FeO tot. Parenthetically, this comparative behaviour shows how sensitive phase equilibria are to small changes in melt starting composition and ambient conditions. This indicates that the approach used here is not one size fits all even though differences in eruptive magmas are relatively small. For T<T inv SiO 2, CaO, FeO tot and K 2 O show a sudden decrease, while Na 2 O and Al 2 O 3 continues to increase as a result of feldspar fractionation. From Fig. 10, MIs hosted in olivine phenocrysts agree with MELTS predicted liquid line of descent, while MI hosted in clinopyroxenes do not. A possible explanation could be post-entrapment diffusive re-equilibration, during cooling; sluggish interface kinetics as T monotonically decreases could contribute to disequilibrium (Qin et al., 1992; Danyushevsky et al., 2000; Cottrell et al., 2002; Michael et al., 2002). If the cooling rate is slow, the diffusive gradient in the crystal may extend to the host magma resulting in re-equilibration between the MI and the magma surrounded the phenocrysts (Gaetani and Watson, 2000). If the cooling rate is fast, such as in the case of scoriae or pumices, post-entrapment re-crystallization could take place as well and the crystallization of the host mineral on the walls of the inclusion modifies the composition of the melt inclusion between entrapment and quenching. We do not have any 18

19 independent evidence of post entrapment re-equilibration in melt inclusions so the cause of the divergence remains open. Good agreement is found between observed and computed clinopyroxene, plagioclase and alkali feldspar compositions (Fig. 11), considering that simulated data are calculated assuming a perfect isobaric fractional crystallization. The closed-system model reproduces the range of observed phenocrysts compositions reasonably well. Although we cannot rule out any involvement of assimilation, there is no indication that this process played a critical petrogenetic role for either the FR or Mi1 system. Small amounts ( several per cent by mass) of assimilation of high Sr crustal contaminant could lead to measured differences in the observed Sr isotopic composition found by D Antonio et al., Changes in properties at T=T inv Significant changes in properties with temperature of melt and magma can be observed in Fig. 12 and 13 respectively for FR and Mi1. All variations in properties become more significant near the invariant temperature T inv, especially for FR. Fig. 12a and 13a shows the variation of melt density with temperature along the liquid line of descent, where the most dramatic change of physical properties for FR and Mi1 occurs at T T inv, because the melt density decreases as a result of a temperature decrease and in residual magma in H 2 O bubbles. Such bubbles would tend to accumulate upwards due to buoyancy effects which would tend to density stratify magma. For example, the density of supercritical H 2 O at T inv for FR (880 C, 0.15 GPa) is ~approximately 300 kg/m3, far smaller than that of melt at the same P and T. The variation of dissolved water in the melt along the liquid line of descent can be observed in Fig. 12b and 13b. For FR, melt saturates with respect to H 2 O at 1108 C at about 4 wt % H 2 O and increases as crystallization occur and heat is extracted. At T inv the H 2 O content jumps from about 4.5 wt% to 5 wt% H 2 O and has a rate of increase of 1 wt% H 2 O per 30 C. For Mi1 the saturation of melt with respect to H 2 O occurs at 800 C at about 8 wt%. Around the invariant interval the value of 19

20 dissolved water jumps from 3.5wt% to 4.3 wt% and increases at the rate of 1.0 wt% per 30 C of cooling. The viscosity of melt as a function of temperature along the crystallization path is shown for both eruptions in Fig. 12c and 13c respectively. For FR the variation of viscosity is similar of what has been observed for the CI (Fowler et al., 2007). Melt viscosity for FR system present a cusped path; a rapid increase with falling of temperature between T liquidus and T inv (due to cooling and the silica enrichment of evolved melt) and then a dramatic drop for T<T inv (due to the increasing concentration of water dissolved in residual melt). As noted on fig. 13c, Mi1 evolution is than FR. In Mi1, the increase of viscosity and dissolved water content as heat is extracted and the temperature falls is more gradual. In Fig. 12d, the volume fraction of water in the magma along the crystallization path for the FR is depicted, where magma has been defined as a homogeneous mixture of oversaturated melt plus bubbles of supercritical fluid (Fowler et al., 2007). The magma density was calculated according to: magma fluid melt 1 ) (1) melt fluid fluid( fluid where fluid is the mass fraction of the fluid phase in the mixture, fluid is the density of exsolved H 2 O and melt is the density of volatile-saturated melt. At T=T inv there is a dramatic increase in volume fraction of water, from about 15% vol to 60% vol just below T inv. The exsolution and expansion of H 2 O provides the mechanical energy that drives explosive volcanic eruptions. According to Cashman et al., (2000), a pyroclastic eruption can occur when the fluid volume fraction exceeds roughly 60-70% by volume at which magma fragmentation occurs. The precise value depends on the distribution of the fluid phase within the magma which might be spatially complex (not homogeneous). The main point is that the volume fraction of the dispersed bubble phase is high enough to lead to a rheological phase transition driven by the change in the identity of the continuous phase -melt to fluid- at least on average. 20

21 Our phase equilibria calculations are consistent with the following picture for FR. Isobaric crystal fractionation of parental basaltic trachyandesitic melt initially containing about 3 wt% H 2 O generates a liquid trajectory in composition space consistent with MI and phenocryst compositional data. In the absence of magma decompression, the crystallization of almost 60% of the original melt and the drastic increase in the volume fraction of exsolved supercritical fluid just below T inv = 880 C, leads to an abrupt increase of the volume of the system and consequent fluid expulsion. This occurs at the same time that the liquid fraction of the system is rapidly decreasing. At this point, a bubble-enriched cap develops at the top of the magma body producing roof hydrofracture and the propagation of volatile-saturated magma filled cracks. The resultant release of pressure during decompression causes further exsolution of fluids from the melt since the solubility of volatiles decreases as pressure is reduced. As the volume fraction of fluid in the magma increases, the magma viscosity also decreases which in turn allows for even more rapid ascent. Via this mechanism of positive feedback the system becomes unstable and an eruption ensues. For Mi1 s magma, the phase equilibria calculations suggest that the system was deeper (~ 12 km depth) and perhaps drier (2 wt% H 2 O) than FR. Both pressure and initial water content imply that the Mi1 magma was further from volatile saturation at depth compared to FR. Unlike the case of FR, for Mi1 simultaneous saturation of plagioclase, alkali feldspar and biotite crystallization took place in a temperature span of ~90 C and not isothermally at T inv. The smaller rate of change of fraction crystallized with temperature naturally leads to less abrupt changes in the melt composition, properties and physical state of the magma. A decrease in melt viscosity (from 10 5 to 10 4 Pa s), coupled with a change in the volume fraction of water from 0.05 to 0.2 and a decrease in melt density nevertheless drove the system towards instability possibly acting as a destabilizing eruption trigger. A prediction of this model is that the Mi1 eruption was less explosive than that of FR. This prediction may be tested by analysis of the volcanic stratigraphy and by granulometric studies on available samples. Poor exposures make this test a difficult one to carry out although one worth trying. 21

22 Timescale for Fondo Riccio and Minopoli 1 magma evolution While phase equilibria modelling can constrain the thermodynamic and transport properties of magmas, the evolutionary timescale cannot be determined without additional considerations. Here we apply a simple thermal model in order to estimate the time interval between the start of fractionation and the eruption in the context of the phase equilibria model. This model can be tested using isotopic data on the various phenocryst phases, although these data do not presently exist. The timescale is estimated by determining the time it takes for sufficient heat to be removed from the magma in order to drive the geochemical evolution from liquidus to eruption temperature. That is, it is assumed that parental melt of volume V (V FR or V MI for FR and Mi1, respectively) and density ρ loses heat at flux rate q and that the total amount of heat that needs to be removed is the difference in enthalpy ( H) between the initial and final states. The fraction of parental melt volume (f m ) that differentiates to form the FR and Mi1 melt compositions and the fraction (α) of that volume that erupts to form FR and Mi1 (respectively V EFR and V EMI ) are linked by the following: V EFR = α f m V FR V EMI = α f m V MI (3a) (3b) The volume of the magma body that crystallizes can be expressed in function of surface area A and a dimensionless constant K that depends on the shape of the magma reservoir, such that A = KV 2/3. The shape of the magma reservoir can be approximated with a cubical, disk-like or spherical volume, for which 7 < K < 5. With these assumptions the timescale can be calculated as: 539 H V ( ) th FR Kq f EFR m 1/ 3 (4a) 540 H V ( ) th MI Kq f EMI m 1/ 3 (4b) 541 The time t since the start of fractionation for each mineral phase is 22

23 t thh (5) Where H is the dimensionless enthalpy and it is function of melt fraction or temperature and is defined: 545 H H liquidus H (T ) H (6) Some parameters such as, H and f m near the solidus are fairly constant and we choose values of 2200 kg/m 3, 1MJ/kg and 0.05 (for Mi1) and 0.1 (for FR). At CF the present day heat flow ( q ) range between 1 and 2.5 W/m 2 (AGIP, 1987; Wohletz et al, 1999; De Lorenzo et al., 2001), as measured at geothermal boreholes in Mofete and San Vito. The fraction (α) of differentiated magma that erupted to form FR and Mi1 eruptive fields can be estimated between 0.5 and 1 (Crisp et al., 1984; White et al., 2006), which we have chosen as the maximum and minimal values. The estimated DRE eruptive volume of FR and Mi1 is 0.16 km 3 and 0.1 km 3 respectively (Di Girolamo et al., 1984) which leads to a timescale τ of 6.5 ± 3.5 ka for FR and 2.5 ± 1.5 ka for Mi1 (Fig. 14ab). The values obtained for τ using the simple thermal model allow us to approximate the timescale for the fractionation process and to give an estimate of the age of each mineral phase. The more evolved compositions of FR MIs and eruptive products can be explained by the longer stationing of the batch magma in the chamber before the eruption, allowing the melt to fractionate up to 60 vol % Conclusions The present study has been conducted with the goal of reconstruction of the pre-eruptive history of the magma bodies that gave rise to the FR and Mi1 deposits. A combination of MI data, thermodynamic and thermal modelling has been brought to bear on this problem. The simulations were carried out rigorously using multicomponent-multiphase phase equilibria tools as embodied in the MELTS algorithm. Both systems were assumed to evolve by fractional crystallization in a 23

24 closed system and computed predictions were compared to observations. MIs in olivine phenocrysts, the first phenocryst to crystallize, evidently represent fossil remnants of the parental magma and were used to represent the starting composition. Parental melt for FR has about 3 wt% H 2 O and evolved by isobaric fractional crystallization at pressure near 0.15 GPa (equivalent to ~8 km of depth) among QFM - QFM+1 oxygen buffer. Calculated phase equilibria along the liquid line of descent show that for P = 0.15 GPa, olivine is the liquidus phase (T liq = 1260 C), followed by clinopyroxene (1110 C), magnetite (1100 C), saturation of water (1070 C) plagioclase, alkali feldspar and biotite (880 C). The calculated oxides trend and composition of phase mineral well agree with observed MIs and mineral assemblage suggesting that FR s system has most likely evolved by closed-system fractional crystallization. At a temperature of 880 C, the magmatic system is subject to a dramatic variation in its physical properties (viscosity, density and water dissolution) as biotite, plagioclase and alkali feldspars start to crystallize. At this temperature, an abrupt decrease in the fraction of melt from 0.5 to 0.1 occurs. The sudden decrease of viscosity and density at this pseudo invariant point temperature and the dramatic change in volume fraction of water from 0.1 to 0.6 is, we speculate, the trigger mechanism for the eruption of FR magma. Mi1 s petrological evolution has been simulated by isobaric fractional crystallization. The starting parental composition based on MI s in olivine suggests a more primitive parent. The system, containing 2 wt% H 2 O, has evolved from pressure of 0.3 GPa and oxygen fugacity values around QFM+1. The crystallization sequence is represented by olivine (T liq = 1300 C), followed by clinopyroxene (T= 1160 C), spinel (T=1070 C), apatite (T=1020 C), plagioclase (T=990 C), biotite (T=960 C) and alkali feldspar (T = 900 C). In the case of Mi1, simulations have not shown invariant temperature behaviour but only a variation of melt fraction (fm) from 0.5 to 0.1 in a temperature span of 90 C (around 990 C), due to the crystallization of alkali feldspars, plagioclase and biotite. A good agreement between observed and calculated mineral compositions suggests that also Mi1 has undergone to a fractional crystallization process even though MIs within later formed clinopyroxene phenocrysts do not appear to represent equilibrium liquids trapped along the liquid 24