1

Understanding Magma Evolution at Campi Flegrei (Campania, Italy) Volcanic 1

Complex Using Melt Inclusions and Phase Equilibria 2

3

Cannatelli C.a,*, Spera F.J.

a, Fedele L.

b, De Vivo B.

c 4

5

a Department of Earth Science and Institute for Crustal Studies, University of California, Santa 6

Barbara, CA 93106 USA 7

b Department of Geosciences Virginia Tech, 4044 Derring Hall, Blacksburg, VA 24061 USA 8

c Dipartimento di Scienze della Terra, Università di Napoli Federico II, 80134 Napoli, Italy 9

10

* Corresponding author: Tel. 1-805-893-8231, Fax: 1-805-893-8649 11

E-mail addresses: claudia@crustal.ucsb.edu (C. Cannatelli), spera@geol.ucsb.edu (F.J. Spera), 12

lfedele@vt.edu (L. Fedele), bdevivo@unina.it (B. De Vivo) 13

14

15

16

17

18

19

20

21

22

23

24

2

Abstract 25

The magmatic evolution of two eruptive episodes at Campi Flegrei (Italy) has been 26

investigated using phase equilibria modeling (MELTS) and data from melt inclusions (MIs) in 27

phenocrysts from the Fondo Riccio (FR, 9.5 ka) and Minopoli 1 (Mi1, 11.1 ka) eruptions. Adopting 28

the Ansatz that isobaric fractional crystallization of a mantle-derived parental magma is the 29

dominant petrogenetic process, major element evolution and corresponding changes in the physical 30

and thermodynamic properties of the magma bodies from which FR and Mi1 magmas were erupted 31

can be tracked. Using olivine hosted MIs as representative of parental melt, the physical conditions 32

and crystallization path have been modeled. Results are compared to observed crystal, whole rock 33

and homogenized MI compositions to evaluate the extent computed phase equilibria can reproduce 34

observations under the imposed conditions. FR parental magma was likely trachyandesitic, 35

approximated by the composition of MIs in olivine (SiO2 = 46.8%, MgO = 9.45 %), which evolved 36

mainly through fractional crystallization at low pressure (P ≈ 0.2 GPa, ≈ 8 km depth), along the 37

QFM±1 oxygen buffer with an initial dissolved H2O content of ~3 wt%. Mi1 parental magma was 38

also trachyandesitic and it is approximated by the chemistry of MIs in olivine (SiO2 = 47.8%, MgO 39

= 9.37%). The estimated mean pressure of crystallization is ≈ 0.3 GPa (≈ 12 km depth), deeper than 40

FR with oxygen fugacity along QFM+1buffer. The initial H2O content of ~ 2 wt% for Mi1 is 41

slightly less than that of FR. Thermodynamic modeling also suggests that mafic parental magma 42

crystallized by about 50% to generate the more evolved (erupted) compositions. MIs in olivine 43

phenocrysts, the first phenocryst to crystallize, evidently represent trapped pristine remnants of the 44

parental magma. MIs within later formed clinopyroxene phenocrysts do not appear to represent 45

equilibrium liquids trapped along the liquid line of descent suggesting that reaction between trapped 46

melt and clinopyroxene may be important or that significant liquid heterogeneity developed by the 47

time clinopyroxene began to crystallize. The relationship between melt fraction and T reveals for 48

FR the presence of a pseudo-invariant temperature, Tinv= 880° at which the fraction of melt 49

decreases abruptly due to simultaneous crystallization of alkali feldspar and plagioclase, eutectic-50

3

like behavior. The melt density, viscosity and dissolved water content change abruptly in a very 51

small temperature interval around Tinv. At this temperature, the volume fraction of exsolved H2O 52

present within magma increases from less than 10% by volume to more than 60 vol % which is of 53

the order of the fragmentation limit of circa 60 vol% for FR differentiated parent melt. In the case 54

of Mi1, simulations do not point to abrupt „invariant temperature behavior‟ but instead melt 55

fraction (fm) varies from 0.5 to 0.2 in a temperature span of 90°C (around 990°C), due to the 56

crystallization of alkali feldspars, plagioclase and biotite. This less „eutectic-like‟ behavior may be 57

due to higher mean crystallization pressure of Mi1 compared to FR. A simple thermal model based 58

on variation of enthalpy of the system along the liquid line of descent allowed us to estimate the 59

duration of the entire differentiation event, suggesting a timescale for FR of 6.5 ± 3.5 kyr and for 60

Mi1 of 2.5±1.5 kyr from the beginning of fractionation until eruption. 61

62

4

1.1 Introduction 63

Campi Flegrei (CF) (Italy) is the most active magmatic system in the Mediterranean region 64

and has exhibited predominantly explosive volcanic activity for more than 300,000 years 65

(Pappalardo et al., 2002). The area is well known for its intense hydrothermal activity, frequent 66

earthquakes and long history of bradyseism including the recent episodes in 1969-1972 and 1982-67

1984. The city of Naples and surroundings, with ~4 million inhabitants, represents one of the most 68

densely populated and volcanically active areas on Earth. The origins of CF‟s explosive volcanism 69

have been the focus of intense research for hundreds of years and is still debated today (Di 70

Girolamo et al, 1984; Rosi and Sbrana, 1987; Barberi et al., 1991; Pappalardo et al., 1999; De Vivo 71

et al., 2001; Rolandi et al., 2003; De Astis et al., 2004; De Vivo and Lima, 2006; Marianelli et al. 72

2006; Bodnar at el., 2007; Di Vito et al., 2008; Lima et al., 2009). 73

Explosive volcanic eruptions constitute a challenge for volcanologists because of their 74

unpredictability; identification of the parameters determining the style of an eruption is of 75

fundamental importance in efforts to understand how explosive volcanoes work. Development of 76

models for volcanic eruption forecasting require information on the pre-eruptive chemical and 77

physical characteristics of the magmatic system (Anderson et al., 2000; Webster et al., 2001; 78

Roggensack et al., 2001; De Vivo et al., 2005; Metrich and Wallace 2008; Moore 2008). In 79

particular the pre-eruptive composition of the magma before the eruption, including its dissolved 80

volatile content, is of critical importance because composition exerts a fundamental control of 81

magma properties and hence the style of eruptive events (Anderson, 1976; Burnham, 1979; De Vivo 82

et al., 2005). The exsolution and expansion of volatiles (especially H2O) provides the mechanical 83

energy that drives explosive volcanic eruptions. The physical properties of magmas, such as density 84

and viscosity, (Lange 1994; Ochs and Lange, 1999; Spera et al, 2000) along with the pre-eruptive 85

phase equilibria (Moore and Carmichael, 1998) are strongly influenced by the dissolution of 86

volatiles in magma and affect the volcanic style of a magmatic system (Sparks et al., 1994). 87

5

Melt inclusions (MIs) are a powerful tool to investigate the pre-eruptive magma composition 88

since they potentially retain the pristine composition of the magma at the time of trapping (Roedder 89

1984). The original volatile content of magma can be estimated by analyzing MIs contained in 90

phenocrysts (Anderson, 1974; Clocchiatti, 1975; Roedder, 1979; Belkin et al., 1985; Sobolev, 1990; 91

Lowenstern, 1994; Anderson, 2003; De Vivo and Bodnar, 2003; Wallace, 2005). Moreover, MIs 92

provide information concerning crystallization and mixing histories of magmas and also the 93

conditions of primary melt generation and extraction (Roedder, 1984; Carroll and Holloway, 1994; 94

Lowenstern, 1994; Sobolev, 1996; Danyushevsky et al., 2000; Frezzotti, 2001). 95

In the past two decades great of effort have been devoted to the description of the processes 96

that drive the evolution of sub-surface magmas at Campi Flegrei as well as the eruptions 97

themselves. In particular, some authors (Civetta et al., 1997; Pappalardo et al. 2002; Tonarini et al. 98

2004; Roach, 2005; D‟Antonio et al., 2007; Arienzo et al., 2009; Tonarini et al., 2009) have shown 99

that fractional crystallization, magma mixing and perhaps wallrock assimilation also play roles in 100

describing the evolution of CF. Indeed sorting out which of these and possibly other mechanisms is 101

most important is a significant part of petrologic research on the evolution of crustal magma bodies. 102

In the present work we examine the origin of magma erupted during the Fondo Riccio, FR (9.5 ka) 103

and Minopoli 1, Mi1 (10.3 ka) volcanic episodes by deriving constraints imposed from phase 104

equilibria embodied in the MELTS thermodynamic model (Ghiorso and Sack, 1995), from 105

phenocryst and glass compositions and from an analysis of MIs found in phenocrysts. Using olivine 106

hosted MIs as representative of parental melt that generated the eruptive products of FR and Mi1, 107

estimates of the pressure, temperature, oxygen buffer, density and viscosity can be made assuming 108

isobaric fractional crystallization was the dominant process of geochemical evolution. Although it 109

would be easy to perform polybaric crystal fractionation (indeed other paths could be chosen) the 110

approach here is to adopt the very simplest case and compare detailed predictions to observations. 111

The deviations from the model and observation then put some constraints on the importance of 112

other processes of petrologic evolution. An important aspect of our findings is the identification of a 113

6

pseudo-invariant temperature (Tinv) along the liquid line of descent. At this temperature the system 114

undergoes dramatic changes in crystallinity, melt composition including volatile content, viscosity, 115

and density. The net effect of these changes is to drive the system towards dynamic instability, 116

which we speculate is the trigger mechanism for the eruptions. A simple thermal model based on 117

the variation of enthalpy of the system along the liquid line of descent is also presented to estimate 118

the timescale between the start of significant crystallization and the time of eruption. 119

120

2.1 Volcanological background 121

Campi Flegrei Volcanic District (CFVD) is a large volcanic complex (~ 200 km2) located west of 122

the city of Naples, Italy (Fig.1). Multiple eruptions have occurred in this area in the last 300 ka 123

(Pappalardo et al., 2002), as well as intense hydrothermal activity, bradyseismic events and frequent 124

earthquakes. The major eruption occurring in the CFVD is the 15 ka Neapolitan Yellow Tuff (NYT) 125

(Deino et al., 2004). The origin of the Campanian Ignimbrite (CI) (39 ka) is controversial: for some 126

authors (Rosi and Sbrana, 1987; Orsi et al., 1996) this eruption occurred in the CFVD; other authors 127

(De Vivo et al., 2001; Rolandi et al., 2003) suggest that the CI originated from fractures activated 128

along the neotectonic Apennine fault system parallel to the Tyrrhenian coastline. They argue that 129

eruptions from >300 ka to 19 ka are not confined to a unique volcanic center or isolated vent system 130

in CF as suggested by Rosi and Sbrana, 1987 and Orsi et al., 1996. De Vivo et al (2001) and 131

Rolandi et al., (2003) argued that only the Neapolitan Yellow Tuff (NYT) (15 ka, Deino et al., 132

2004) erupted from vents within CF, whereas the CI (39 ka, DeVivo et al., 2001) has a much wider 133

source and dispersal area. 134

According to Pappalardo et al. (2002), the interval between the CI and NYT eruptions is 135

characterized by a number of small magnitude volcanic events. Since the NYT eruption, the 136

margins of the region have been the site of at least 65 eruptions, divided in three periods of activity. 137

Eruptions were separated by quiescent periods marked by two widespread paleosols (Di Vito et al., 138

7

1999). The last eruption in 1538 A.D. formed the Monte Nuovo cone (Di Vito et al., 1987) after 3.4 139

ka of dormancy. 140

In this paper we analyze the Fondo Riccio (FR) and Minopoli 1 (Mi1) eruptive products in an 141

effort to deduce their petrogenesis. The FR eruption was explosive with a strombolian character and 142

occurred at 10.3 - 9.5 kyr (D‟Antonio et al., 1999) from an eruptive centre on the western side of 143

the Gauro volcano, near the centre of the Phlegrean caldera (Fig 1). The eruptive deposits are 144

limited to the vent area and lie above the Paleosol A and below the Montagna Spaccata Tephra. The 145

eruptive products consist of fallout deposits composed of very coarse scoria beds with subordinate 146

coarse ash beds (Di Vito et al., 1999). 147

According to Di Vito et al. (1999), the earlier Mi1 eruption occurred 10.3 - 11.1 ka and was 148

strombolian with subordinate phreatomagmatic phases, while Di Girolamo et al. (1984), based on 149

the degree of dispersal of Mi1‟s products, define this eruption as sub-Plinian. The deposits are 150

limited to the vent area formed by scoriae horizons with a composition varying from latitic to alkali-151

trachytic. The eruptive products are composed of alternating pumice lapilli fallout and mainly 152

massive ash fallout beds and, subordinately, cross laminated ash surge beds, rich in accretionary 153

lapilli (Di Vito et al., 1999). Evidently, the Mi1 eruption had a stronger phreatomagmatic 154

component than the closely related FR eruption based on observed stratigraphy. 155

156

3.1 Sample description and analytical technique 157

The locations of the samples utilized in this study are indicated in Figure 1. Here we give 158

petrographic and mineralogical descriptions of the samples and describe the methods used to 159

perform the analyses. 160

161

3.1.1 Petrography and chemical composition of Fondo Riccio 162

8

For FR, CF-FR-C1 was collected at the top of the stratigraphic column and is a well-163

vesciculated scoriae containing approximately 20% by volume of phenocrysts. The phenocrysts 164

include olivine, clinopyroxene, spinel (magnetite), biotite, alkali feldspar and plagioclase. Biotite 165

occurs as large crystals (typical size ~ 2-3 mm), while apatite phenocrysts occur as small (~ 0.1 166

mm) acicular needles. Clinopyroxene and feldspar commonly exhibit intergrowth textures, 167

suggesting cotectic crystallization. Olivine, clinopyroxene and plagioclase contain recrystallized 168

MIs, while alkali feldspar phenocrysts contain apatite inclusions. Sample, CF-FR-C2, is a bomb, 169

relatively unvesciculated, containing olivine, clinopyroxene, apatite, spinel, biotite, alkali feldspar 170

and plagioclase. Olivine, clinopyroxene and alkali feldspar phenocrysts contain recrystallized MIs. 171

Petrochemically, both samples are porphyritic latite with ~ 20% phenocrysts, with clinopyroxene 172

and plagioclase often found in glomeroporphyritic clots; clinopyroxene and plagioclase also occur 173

as microlites in the groundmass. 174

In the FR samples, olivine phenocrysts range Fo84- 87, and pyroxene lies in the diopside-salite 175

field on the pyroxene quadrilateral, with Wo44-47 and Fs6-15 (Table 1). Based on microprobe 176

analyses, alkali feldspars in FR present a unimodal distribution with Or component of ~ 79 to 88. 177

Plagioclase crystals are zoned with An 72-98 (Table 2). 178

179

3.1.2 Petrography and chemical composition of Minopoli 1 180

For Mi1, CF-MI1-C1 was collected in the Casalesio area (Fig 1), at the base of the deposit. 181

The sample is greyish-black scoriae, of trachybasalt composition containing ~ 20% phenocrysts of 182

olivine, clinopyroxene, plagioclase, alkali feldspar, spinel (magnetite), apatite and biotite. Olivine 183

phenocrysts are weakly to unzoned with average Fo content ~ 78, while pyroxenes present Wo 184

values between 45 and 48 and Fs between 6 and 16 (Table 3). Based on microprobe analyses, alkali 185

feldspars in Minopoli 1 present a unimodal distribution of Or values which ranges from ~ 75 to 80. 186

Alkali feldspars exhibit zonation, with potassic cores. Plagioclase crystals are highly zoned 187

9

presenting a bimodal distribution with a range from ~ 54 to 87 with peaks at 54 and 83 based on 188

about 50 grains (Table 4). 189

190

3.1.3 Melt Inclusions description 191

The MI‟s present in both FR and M1 are generally devitrified and partially recrystallized, 192

present a bubble (shrinkage ± exsolution of volatiles) and daughter minerals (generally apatite and 193

oxides). MIs generally have elongated ellipsoidal shapes and range from 30 to 80 µm (most 194

between 20 and 50 µm). In order to be analyzed, MIs needed to be re-heated to a homogenous glass. 195

Detailed descriptions of MIs reheating procedures, sample preparation and analytical methods are in 196

Cannatelli et al., 2007. 197

198

3.1.4 Analytical methods 199

Major and minor elements analyses of phenocrysts were performed in the Department of 200

Earth Science at UCSB using a Cameca SX-50 electron microprobe equipped with five wavelength 201

dispersive spectrometers. Phenocrysts analyses were performed using a 1µm focused beam at 15 202

keV accelerating voltage and a beam current of 15nA. Uncertainty of analyses is around 1% 203

(relative) for most elements. Quantitative electron microprobe analyses (EMPA) on phenocrysts and 204

MIs were performed at Virginia Tech and at University of Rome “La Sapienza” (IGAG-CNR, 205

Rome, Italy) on a Cameca SX-50 equipped with four wavelength dispersive spectrometers. The 206

analytical scheme chosen for MIs is described in Cannatelli et al., 2007 and reference therein. 207

208

4.1 Phase equilibria modeling 209

4.1.1 Procedures to select the parental melt composition 210

Phase equilibria modeling has been carried out using the software MELTS, a thermodynamic 211

model of crystal-liquid equilibria. The MELTS algorithm is based on classical equilibrium 212

10

thermodynamics and has been object of extensive reviews (Ghiorso and Sack, 1995, Asimow and 213

Ghiorso, 1998). The use of MELTS to reconstruct the crystallization path of a magma requires 214

specification of initial conditions, including 1) the initial state of the system (parental melt 215

composition including H2O content, starting temperature and pressure, and oxygen fugacity) and 2) 216

constraints under which the magmatic evolution proceeds (open or closed system, fractional or 217

equilibrium crystallization, minimization of appropriate thermodynamic potential based on imposed 218

constraints). In this work we investigate isobaric crystallization scenarios and explore both 219

equilibrium and fractional crystallization scenarios. The search of parental melt composition starts 220

with the assumption that MIs within phenocryst phases can be related to a unique parental melt 221

during cotectic (olivine +clinopyroxene) crystallization. The graphical method developed by 222

Watson (1976) is used to test the hypothesis that MIs are primary or nearly so. MIs composition(s) 223

of interest are further culled by selecting ones that exhibit the lowest concentrations of incompatible 224

trace elements and highest MgO contents as input for the phase equilibria calculations. 225

In the case of Fondo Riccio, 7 MIs were selected, hosted in olivine and pyroxene and have 226

been plotted on a CaO-MgO-Al2O3 coordinates, as described by Watson (1976). The intersection I 227

(Figure 2a) of olivine and clinopyroxene fractionation lines is in the field occupied by FR-C1-o6 228

M1, a MIs hosted in olivine O6. This MI represents the predicted composition of the melt at the 229

cotectic point, where olivine and clinopyroxene crystallize simultaneously, so it is reasonable to 230

hypothesize that the Parental Melt (PM) composition should be more primitive than FR-C1-o6 M1. 231

The MIs FR-C1-o2 M1 (9.45 wt% MgO), and FR-C1-o1 M1 (8.05 wt % MgO) possess high MgO 232

contents and the lowest concentration of incompatible trace elements and are consequently 233

considered the best candidates to represent the parental melt. We carried out phase equilibria 234

calculations using FR-C1-o1 M1 (not shown) and FR-C1-o2 M1 and differences were small; based 235

on this we decided to select the one with the highest MgO content. 236

In the case of Mi1, by applying the Watson graphical method we found that Mi1-C1-P8 M1, a MI 237

hosted in the clinopyroxene P8 (fig 2b) represents the composition of the melt at the cotectic point. 238

11

We selected the parental melt composition choosing the MI with the highest MgO content and 239

lowest incompatible trace element concentrations as an approximation to the PM. The MI that best 240

fit the criteria and was closest to Mi1-C1-P8 M1 in Fig. 2b was hosted in olivine o5 with a MgO 241

content of 9.37 wt%, and values of Ce, and Nd of 69 and 61ppm. It is probable that MIs in olivine 242

can undergo some re-equilibration with the host (Danyushevsky et al., 2000; Kress and Ghiorso, 243

2004). However in our case the MELTS results agree very well with the compositions for the MIs 244

in olivine for both FR and Mi1 samples. Our interpretation of these relations is that that post 245

entrapment changes for these MIs are small to negligible. We conclude that the method espoused 35 246

years ago by Watson is indeed useful and that by careful use of MIs one can at least in this case 247

estimate the parental melt composition reasonably well. 248

249

4.1.2 Phase equilibria: constraints and limitations 250

To reconstruct the magmatic evolution the initial state of the system, devolatilized PM 251

composition, dissolved H2O content of PM, initial temperature, pressure, and oxygen buffer are 252

specified. Here we present results of closed system isobaric fractional crystallization where the 253

Gibbs energy is the appropriate thermodynamic potential to be minimized. We have adopted these 254

constraints as an Ansatz to be tested by the closeness of the computed results to observations. These 255

runs clearly show the effects of varying pressure, fO2 and the initial water content of the parental 256

melt on the liquid line of descent and on the composition and abundance of all crystalline phases 257

and the temperature at which melt becomes water saturated. After setting P, fO2 and dissolved H2O 258

content, we compare predicted phase and melt compositions to those observed in order to determine 259

the range of physical conditions leading up to eruption for FR and Mi1. We selected the “best case” 260

based on correspondence between mineralogical and geochemical data and the phase equilibria 261

calculations. Calculations were rejected when the deviation between observation and model was 262

deemed too large. Although the degree of „closeness‟ could be quantified by, for example, using a 263

Euclidean norm criterion comparing the predicted oxide composition of a phase to its observed 264

12

values, we believe such a procedure is a premature at the present time. Instead, we prefer to rely on 265

reasonable judgment predicated on the assumption that an experienced petrologist will be able to 266

spot a poor solution as one that provides no new insight into the petrogenesis of the system and on 267

what may have triggered the eruptions. One must keep in mind the assumptions of the method and 268

the realities of Nature. For example, the calculation assumes perfect fractional crystallization. 269

However, in situations where crystals are removed from liquid by some physical process driven by 270

gravity (e.g., crystal settling/floatation) or deviatoric stress (e.g., kneading, melt percolation, filter 271

pressing, see Kohlstedt and Holtzman, 2009), there will always be some reaction between earlier 272

formed crystals and ambient liquid. Similarly, the calculation assumes there is a single parental 273

composition from which all differentiated liquids develop. It is easy to imagine that compositional 274

heterogeneities would be present a priori even if convective mixing was reasonably efficient. 275

Finally, the calculation assumes that crystallization is isobaric, exactly. The approximate nature of 276

this assumption should be clear to anyone who ever mapped plutonic in rugged terrain. The point of 277

performing phase equilibria calculations using an imperfect thermodynamic model (no 278

thermodynamic model is perfect) with constraints that are clearly approximate is to evaluate the 279

overall reasonability of the proposed scenario. If, for example, crystallization is grossly polybaric, 280

then no isobaric model will come close to reproducing observed phase compositions, abundances 281

and glass (melt) compositions. One could then perform a constrained polybaric simulation and ask 282

if that procedure produces better agreement. If assimilation plays an important part of the 283

petrogenesis, then no closed system phase equilibria model will produce satisfactory 284

correspondence to observations and one would seek to explore alternative petrogenetic models 285

quantitatively involving assimilation and the mixing of melts or magmas of differing composition 286

and temperature. 287

In this study we find that isobaric closed system fractional crystallization at low pressure 288

produces results that bear a close (but not perfect) correspondence to observed relations and that the 289

13

implications of the calculation suggest a causative link between crystallization and the eruptive 290

episode that generated the two small volcanic deposits of the FR and Mi1 (see below). 291

292

4.2 Fondo Riccio 293

The initial water content in the parental melt has been estimated starting from the values 294

obtained for MIs by SIMS analyses. FR‟s MIs belong to two different populations of inclusions, 295

one with water contents ranging between 1 and 4 wt% and the other with water values around 6 296

wt%. As starting water content we tested values ranging between 1 and 5 wt%, but from 297

petrographic observations values of H2O >3wt% were discarded because of the high water 298

saturation temperature. For example, in the case of 4wt% H2O the temperature of water saturation 299

was 1070°C at 0.2 GPa (depth ~ 6 km). At this temperature the system is saturated in water and 300

crystallizing mineral phases such as clinopyroxene, plagioclase and alkali feldspar should trap fluid 301

inclusions during the cooling process. There is no petrographic evidence of fluid inclusions hosted 302

in these phases in the samples studied here. In the cases of H2O < 2 wt%, each run generated a 303

rhombohedral oxide phase (illmenite) at low melt fractions, inconsistent with the phase assemblage 304

observed. Although not shown, calculated runs with initial water content in the PM less than 2 wt% 305

and greater than 4 wt% did not predict the phase assemblage observed in the FR. We therefore 306

conclude that initial water content in the PM around 3 wt% is the most realistic case for the FR 307

eruptive system. Although we acknowledge that this is a judgment, we believe it to be the best 308

estimate based on the congruence between calculation and what is observed in the natural samples 309

studied in the laboratory. 310

The majority of the runs were made isobarically and for FR at P < 0.3 GPa; at greater P the 311

presence of predicted minerals such as garnet or muscovite is not compatible with the FR 312

phenocryst assemblage. To understand better the effect of changing pressure, we compared MELTS 313

generated TAS diagrams calculated at a fixed fO2 = QFM+1, QFM and P = 0.1, 0.15, 0.2 and 0.3 314

GPa. For the case of fO2 = QFM and QFM+1 good agreement between phase equilibria (MELTS) 315

14

with FR‟s data (see Fig. 3). The best case scenario of oxygen fugacity for FR was chosen for P ≈ 316

0.2 GPa, corresponding to ~8 km depth, and compatible with recent studies by Zollo et al., 2008 317

suggesting that a hypothetical magma body at Campi Flegrei is at least 7.5 km deep. 318

From petrographic investigation we found the presence of spinel (in the form of magnetite solid 319

solution) in olivine and clinopyroxene, but not in plagioclase and feldspars. Biotite is also present. 320

We compared several MELTS generated mineral distribution diagrams with petrographic 321

observations and found best agreement is reached for fO2 between QFM-1 and QFM+1. We also 322

noticed, as expected, the strong dependence of the iron-bearing phases on the variation of oxygen 323

fugacity. For example, when we consider the case of FR with initial water content of 2 wt%, an 324

increase in the oxygen fugacity from QFM-2 to QFM+2, stabilizes spinel at higher temperature, 325

while not affecting the crystallization temperature of clinopyroxenes and feldspars (Fig. 4). The 326

stabilization of spinel at higher temperatures corresponds to a decrease of FeOtot and increase of 327

SiO2 content in the melt. The inconsistency between observed mineral assemblage and MELTS 328

generated mineral distribution has lead us to discard oxygen fugacity extreme values of QFM-2, 329

QFM-1 and QFM+2. 330

In summary, the physical conditions that produce the closest correspondence between the 331

model and observation is fractional crystallization of a parental melt of (anhydrous) composition 332

(given in Table 5) plus 3 wt % H2O added at 0.15 GPa and oxygen fugacity around the QFM buffer. 333

334

4.3 Minopoli 1 335

Water contents of MIs from the Mi1 eruptive products were measured by SIMS and range 336

from 1 to 4 wt% (Cannatelli et al., 2007). The effect of varying the initial water concentration in the 337

parental melt was examined in the Mi1 case through isobaric fractional crystallization as for FR. 338

Petrographic studies of Mi1‟s thin sections reveal the presence of large (1-2 mm) biotite crystals. 339

The presence of such crystals implies initial water contents greater than 2 wt%. Therefore 340

simulations obtained by setting the water content less than 2wt% were discarded, regardless of 341

15

oxygen fugacity and pressure values. Furthermore, in the case of H2O > 2wt% we observed a lack 342

of intersection between the MELTS generated oxides trends and the real data field of Mi1. In 343

particular, values of water content greater of 3 wt% were discarded for fO2 = QFM ≥ QFM+2 and 344

pressure greater than 0.3 GPa, because of the predicted presence of garnet and leucite, inconsistent 345

with the observed assemblage. Values of water greater than 4 wt% were discarded because of the 346

high water saturation temperature (T ~ 1080°C) which would result in the presence of fluid 347

inclusions in the phenocrysts of Mi1 sample, not observed in Mi1. The initial water content of the 348

parental melt for Mi1 estimated is therefore around 2 wt% a bit lower than that FR using the same 349

methods. 350

Several simulations were carried out using a fixed value of initial water content of 2-3wt%, 351

and varying the pressure and the oxygen fugacity. Many runs were discarded because of mismatch 352

between observed and predicted phases, such in the cases of fO2 > QFM or P ≤ 0.1GPa. A small 353

decrease in oxygen fugacity leads to a decrease of spinel stabilization temperature of almost 100°C 354

and a longer crystallization interval for feldspars with a consequent greater generated mass of 355

feldspars in the mineral assemblage. Comparisons among feldspars plotting model results and 356

observations on ternary diagrams (An-Ab-Or) and spinel diagrams (FeO-Fe2O3-TiO2) were studied 357

in order to establish the best fit. In general higher pressures better match observed phases (Fig. 5). 358

In particular, for fO2 = QFM+1 we obtain a good fit for spinel and feldspar compositions to 359

observed M1 phenocryst (see Table 6), especially at P=0.3GPa and water content of 2 wt%. The 360

best case chosen from the several M1 simulations is represented by a parental melt of (anhydrous) 361

composition (Table 7) at pressure P ~ 0.3 GPa (~12 km depth), water content of 2 wt % along the 362

QFM+1 oxygen buffer. 363

364

5.1 Results 365

5.1.1 Fondo Riccio 366

16

We present results for FR for isobaric fractional crystallization of the estimated parental 367

composition. In fact, we have used a number of possible parental compositions and although small 368

differences in results are obtained, the salient features are not significantly affected. The parental 369

melt composition of FR-C1-o2 M1 with an initial water content of 3 wt% is used to generate the 370

results below. The fractional crystallization path along the QFM to QFM+1 oxygen buffer at 0.15 371

GPa has been computed. MELTS correctly predicts the mineral phases observed. Olivine is the 372

liquidus (T= 1260°C) phase, followed by clinopyroxene, magnetite, H2O, plagioclase, Alkali 373

feldspar and biotite at 1110°C, 1100°C, 1070°C and 880°C respectively. Mineral distribution, 374

abundances and temperature at which water saturates are shown in Fig. 3. It is interesting to note 375

the abrupt change in melt composition around 880°C due to a simultaneous crystallization of alkali 376

feldspar, plagioclase and biotite. We can define this temperature as a pseudo-invariant point 377

temperature (Fowler et al., 2007). At this fixed temperature, a major change in melt fraction (fm), 378

from 0.5 to 0.1 and melt composition occurs (Fig. 6) as crystals precipitate from the melt due to 379

extraction of heat (enthalpy). The properties of the melt (density and shear viscosity) and of magma 380

(density, volume fraction of bubbles, shear viscosity) change dramatically around this temperature 381

(see below). 382

The growth of alkali feldspars and plagioclase dominates the crystallization path at T ~ 880°C 383

and below. In fig. 4 crystallization patterns for fO2 = QFM and QFM+1 are portrayed. 384

Concentrations of SiO2, K2O, Na2O and Al2O3 initially increase with decreasing MgO due to the 385

crystallization of olivine and continue to increase as clinopyroxene, spinel and apatite crystallize 386

(Fig. 7a-f). The increase of CaO concentration ends when melt becomes saturated in clinopyroxene 387

and then decreases slowly with cooling. FeOtot

concentrations slightly decrease in the early stages 388

of crystallization and decrease abruptly when spinel joins the already fractionated phase minerals. 389

Results for QFM and QFM+1 are quite similar for Al2O3, K2O and Na2O while for SiO2, FeOtot

and 390

CaO we can observe a more close approximation to the observed trends for fO2=QFM+1. 391

17

At T=Tinv there is a rapid change in the variation diagram trajectories of circa 2 wt% for SiO2 and 392

Al2O3, 1 wt% for K2O and 0.5 wt% for CaO and Na2O. For T<Tinv SiO2, CaO, Al2O3 and K2O show 393

a sudden decrease, while Na2O continues to increase as a result of feldspar fractionation. These 394

compositional changes at T=Tinv are associated to a change in the physical properties of both melt 395

and magma with significant consequences for eruption probability and dynamics. 396

As noted on Fig. 7, MI‟s hosted in olivine and clinopyroxene agree well with the predicted 397

liquid line of descent making melt inclusions, especially hosted in olivine. There is a good 398

agreement between observed and simulated clinopyroxene and olivine compositions (Fig. 8) 399

remembering that calculated values assume perfect fractional crystallization. Alkali feldspar trends 400

compare favourably; predicted plagioclase becomes more sodic than observed values near the 401

solidus presumably related to the breakdown of the assumption of perfect fractional crystallization 402

near the solidus where the separation of melt from surrounding solids is slowed, melt percolation 403

through a mush being much slower than crystal settling in a crystal-poor magma. We tested two 404

cases of fractionation: (1) All crystals and exsolved H2O (water bubbles) are removed immediately 405

upon saturation and (2) only precipitated crystals are fractionated; water bubbles remain suspended 406

in the melt. The best agreement was found when both precipitated solids and exsolved H2O were 407

removed in „perfect‟ fractional crystallization. The spinel ternary diagram also shows good 408

agreement between MELTS predictions and spinel compositions from FR samples (Table 5). 409

410

5.1.2 Minopoli 1 411

Based on the assumption that the most realistic parental melt has a composition of Mi1-C1-o5- 412

with water content of 2-3 wt%, we present results for calculations with oxygen buffer set at QFM+1 413

and pressure at 0.2-0.3 GPa. In Fig. 9 we can see the mineral distribution along the crystallization 414

path for the case P= 0.3 GPa, 2wt% H2O and fO2= QFM+1. 415

18

The liquidus phase is olivine at T= 1300°C, followed by clinopyroxene (T= 1160°C), spinel 416

(T=1070°C) and apatite (T=1020°C). At 990°C, 960°C and 900°C respectively plagioclase, biotite 417

and alkali feldspar join the mineral assemblage, dominating the crystallization path. 418

From Fig. 10a-f, concentrations of SiO2, Al2O3, K2O and Na2O increase with decreasing MgO 419

during the crystallization of olivine and then continue to increase as clinopyroxene, apatite and 420

spinel crystallize. The increase of CaO ends when clinopyroxene begins crystallization and then 421

decreases slowly with cooling. FeOtot

slightly decreases in the early stages of crystallization, then 422

remains constant and decreases abruptly only when spinel begins fractionation. In the case of Mi1 423

an abrupt change in melt composition noted in FR is not evident; instead, in a temperature span of 424

about 80°C (around T= 990°C) there is a change of fm from 0.5 to 0.2, due mainly to the 425

crystallization of feldspar. At T=990°C, there are changes in the calculated oxides trends of about 3 426

wt% for SiO2, CaO and K2O, 2wt% for Al2O3 and 1 wt% for Na2O and FeOtot

. Parenthetically, this 427

comparative behaviour shows how sensitive phase equilibria are to small changes in melt starting 428

composition and ambient conditions. This indicates that the approach used here is not „one size fits 429

all‟ even though differences in eruptive magmas are relatively small. 430

For T<Tinv SiO2, CaO, FeOtot

and K2O show a sudden decrease, while Na2O and Al2O3 continues to 431

increase as a result of feldspar fractionation. From Fig. 10, MIs hosted in olivine phenocrysts agree 432

with MELTS predicted liquid line of descent, while MI hosted in clinopyroxenes do not. A possible 433

explanation could be post-entrapment diffusive re-equilibration, during cooling; sluggish interface 434

kinetics as T monotonically decreases could contribute to disequilibrium (Qin et al., 1992; 435

Danyushevsky et al., 2000; Cottrell et al., 2002; Michael et al., 2002). If the cooling rate is slow, the 436

diffusive gradient in the crystal may extend to the host magma resulting in re-equilibration between 437

the MI and the magma surrounded the phenocrysts (Gaetani and Watson, 2000). If the cooling rate 438

is fast, such as in the case of scoriae or pumices, post-entrapment re-crystallization could take place 439

as well and the crystallization of the host mineral on the walls of the inclusion modifies the 440

composition of the melt inclusion between entrapment and quenching. We do not have any 441

19

independent evidence of post entrapment re-equilibration in melt inclusions so the cause of the 442

divergence remains open. 443

Good agreement is found between observed and computed clinopyroxene, plagioclase and 444

alkali feldspar compositions (Fig. 11), considering that simulated data are calculated assuming a 445

perfect isobaric fractional crystallization. The closed-system model reproduces the range of 446

observed phenocrysts compositions reasonably well. Although we cannot rule out any involvement 447

of assimilation, there is no indication that this process played a critical petrogenetic role for either 448

the FR or Mi1 system. Small amounts ( several per cent by mass) of assimilation of high Sr crustal 449

contaminant could lead to measured differences in the observed Sr isotopic composition found by 450

D´Antonio et al., 2007. 451

452

5.1.3 Changes in properties at T=Tinv 453

Significant changes in properties with temperature of melt and magma can be observed in Fig. 454

12 and 13 respectively for FR and Mi1. All variations in properties become more significant near 455

the invariant temperature Tinv, especially for FR. Fig. 12a and 13a shows the variation of melt 456

density with temperature along the liquid line of descent, where the most dramatic change of 457

physical properties for FR and Mi1 occurs at T ≤ Tinv, because the melt density decreases as a result 458

of a temperature decrease and in residual magma in H2O bubbles. Such bubbles would tend to 459

accumulate upwards due to buoyancy effects which would tend to density stratify magma. For 460

example, the density of supercritical H2O at Tinv for FR (880 C, 0.15 GPa) is ~approximately 300 461

kg/m3, far smaller than that of melt at the same P and T. 462

The variation of dissolved water in the melt along the liquid line of descent can be observed in 463

Fig. 12b and 13b. For FR, melt saturates with respect to H2O at 1108°C at about 4 wt % H2O and 464

increases as crystallization occur and heat is extracted. At Tinv the H2O content jumps from about 465

4.5 wt% to 5 wt% H2O and has a rate of increase of 1 wt% H2O per 30°C. For Mi1 the saturation of 466

melt with respect to H2O occurs at 800°C at about 8 wt%. Around the invariant interval the value of 467

20

dissolved water jumps from 3.5wt% to 4.3 wt% and increases at the rate of 1.0 wt% per 30°C of 468

cooling. 469

The viscosity of melt as a function of temperature along the crystallization path is shown for 470

both eruptions in Fig. 12c and 13c respectively. For FR the variation of viscosity is similar of what 471

has been observed for the CI (Fowler et al., 2007). Melt viscosity for FR system present a cusped 472

path; a rapid increase with falling of temperature between Tliquidus and Tinv (due to cooling and the 473

silica enrichment of evolved melt) and then a dramatic drop for T<Tinv (due to the increasing 474

concentration of water dissolved in residual melt). As noted on fig. 13c, Mi1 evolution is than FR. 475

In Mi1, the increase of viscosity and dissolved water content as heat is extracted and the 476

temperature falls is more gradual. 477

In Fig. 12d, the volume fraction of water in the magma along the crystallization path for the 478

FR is depicted, where magma has been defined as a homogeneous mixture of oversaturated melt 479

plus bubbles of supercritical fluid (Fowler et al., 2007). The magma density was calculated 480

according to: 481

)1( fluidfluidfluidmelt

meltfluid

magma

(1) 482

where fluid is the mass fraction of the fluid phase in the mixture, fluid is the density of exsolved 483

H2O and melt is the density of volatile-saturated melt. At T=Tinv there is a dramatic increase in 484

volume fraction of water, from about 15% vol to 60% vol just below Tinv. The exsolution and 485

expansion of H2O provides the mechanical energy that drives explosive volcanic eruptions. 486

According to Cashman et al., (2000), a pyroclastic eruption can occur when the fluid volume 487

fraction exceeds roughly 60-70% by volume at which magma fragmentation occurs. The precise 488

value depends on the distribution of the fluid phase within the magma which might be spatially 489

complex (not homogeneous). The main point is that the volume fraction of the dispersed bubble 490

phase is high enough to lead to a rheological „phase transition‟ driven by the change in the identity 491

of the continuous phase -melt to fluid- at least on average. 492

21

Our phase equilibria calculations are consistent with the following picture for FR. Isobaric 493

crystal fractionation of parental basaltic trachyandesitic melt initially containing about 3 wt% H2O 494

generates a liquid trajectory in composition space consistent with MI and phenocryst compositional 495

data. In the absence of magma decompression, the crystallization of almost 60% of the original melt 496

and the drastic increase in the volume fraction of exsolved supercritical fluid just below Tinv= 497

880°C, leads to an abrupt increase of the volume of the system and consequent fluid expulsion. This 498

occurs at the same time that the liquid fraction of the system is rapidly decreasing. At this point, a 499

bubble-enriched cap develops at the top of the magma body producing roof hydrofracture and the 500

propagation of volatile-saturated magma–filled cracks. The resultant release of pressure during 501

decompression causes further exsolution of fluids from the melt since the solubility of volatiles 502

decreases as pressure is reduced. As the volume fraction of fluid in the magma increases, the 503

magma viscosity also decreases which in turn allows for even more rapid ascent. Via this 504

mechanism of positive feedback the system becomes unstable and an eruption ensues. 505

For Mi1‟s magma, the phase equilibria calculations suggest that the system was deeper (~ 12 506

km depth) and perhaps drier (2 wt% H2O) than FR. Both pressure and initial water content imply 507

that the Mi1 magma was further from volatile saturation at depth compared to FR. Unlike the case 508

of FR, for Mi1 simultaneous saturation of plagioclase, alkali feldspar and biotite crystallization took 509

place in a temperature span of ~90°C and not isothermally at Tinv. The smaller rate of change of 510

fraction crystallized with temperature naturally leads to less abrupt changes in the melt composition, 511

properties and physical state of the magma. A decrease in melt viscosity (from 105 to 10

4 Pa s), 512

coupled with a change in the volume fraction of water from 0.05 to 0.2 and a decrease in melt 513

density nevertheless drove the system towards instability possibly acting as a destabilizing eruption 514

trigger. A prediction of this model is that the Mi1 eruption was less explosive than that of FR. This 515

prediction may be tested by analysis of the volcanic stratigraphy and by granulometric studies on 516

available samples. Poor exposures make this test a difficult one to carry out although one worth 517

trying. 518

22

519

6.1 Timescale for Fondo Riccio and Minopoli 1 magma evolution 520

While phase equilibria modelling can constrain the thermodynamic and transport properties of 521

magmas, the evolutionary timescale cannot be determined without additional considerations. Here 522

we apply a simple thermal model in order to estimate the time interval between the start of 523

fractionation and the eruption in the context of the phase equilibria model. This model can be tested 524

using isotopic data on the various phenocryst phases, although these data do not presently exist. 525

The timescale is estimated by determining the time it takes for sufficient heat to be removed from 526

the magma in order to drive the geochemical evolution from liquidus to eruption temperature. That 527

is, it is assumed that parental melt of volume V (VFR or VMI for FR and Mi1, respectively) and 528

density ρ loses heat at flux rate q and that the total amount of heat that needs to be removed is the 529

difference in enthalpy (H) between the initial and final states. The fraction of parental melt volume 530

(fm) that differentiates to form the FR and Mi1 melt compositions and the fraction (α) of that 531

volume that erupts to form FR and Mi1 (respectively VEFR and VEMI) are linked by the following: 532

VEFR = α fm VFR (3a) 533

VEMI = α fm VMI (3b) 534

The volume of the magma body that crystallizes can be expressed in function of surface area A and 535

a dimensionless constant K that depends on the shape of the magma reservoir, such that A = KV2/3

. 536

The shape of the magma reservoir can be approximated with a cubical, disk-like or spherical 537

volume, for which 7 < K < 5. With these assumptions the timescale can be calculated as: 538

3/1

)(

m

EFRth

f

V

qK

HFR

(4a) 539

3/1

)(

m

EMIth

f

V

qK

HMI

(4b) 540

The time t since the start of fractionation for each mineral phase is 541

23

Ht th (5) 542

Where H is the dimensionless enthalpy and it is function of melt fraction or temperature and is 543

defined: 544

H H liquidusH (T )

H (6) 545

546

Some parameters such as , H and mf near the solidus are fairly constant and we choose values 547

of 2200 kg/m3, 1MJ/kg and 0.05 (for Mi1) and 0.1 (for FR). At CF the present day heat flow ( q ) 548

range between 1 and 2.5 W/m2

(AGIP, 1987; Wohletz et al, 1999; De Lorenzo et al., 2001), as 549

measured at geothermal boreholes in Mofete and San Vito. The fraction (α) of differentiated magma 550

that erupted to form FR and Mi1 eruptive fields can be estimated between 0.5 and 1 (Crisp et al., 551

1984; White et al., 2006), which we have chosen as the maximum and minimal values. The 552

estimated DRE eruptive volume of FR and Mi1 is 0.16 km3 and 0.1 km

3 respectively (Di Girolamo 553

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. 14a-554

b). The values obtained for τ using the simple thermal model allow us to approximate the timescale 555

for the fractionation process and to give an estimate of the age of each mineral phase. The more 556

evolved compositions of FR MIs and eruptive products can be explained by the longer stationing of 557

the batch magma in the chamber before the eruption, allowing the melt to fractionate up to 60 vol 558

%. 559

560

7. 1 Conclusions 561

The present study has been conducted with the goal of reconstruction of the pre-eruptive history 562

of the magma bodies that gave rise to the FR and Mi1 deposits. A combination of MI data, 563

thermodynamic and thermal modelling has been brought to bear on this problem. The simulations 564

were carried out rigorously using multicomponent-multiphase phase equilibria tools as embodied in 565

the MELTS algorithm. Both systems were assumed to evolve by fractional crystallization in a 566

24

closed system and computed predictions were compared to observations. MIs in olivine 567

phenocrysts, the first phenocryst to crystallize, evidently represent fossil remnants of the parental 568

magma and were used to represent the starting composition. Parental melt for FR has about 3 wt% 569

H2O and evolved by isobaric fractional crystallization at pressure near 0.15 GPa (equivalent to ~8 570

km of depth) among QFM - QFM+1 oxygen buffer. Calculated phase equilibria along the liquid 571

line of descent show that for P = 0.15 GPa, olivine is the liquidus phase (Tliq = 1260°C), followed 572

by clinopyroxene (1110°C), magnetite (1100°C), saturation of water (1070°C) plagioclase, alkali 573

feldspar and biotite (880°C). The calculated oxides trend and composition of phase mineral well 574

agree with observed MIs and mineral assemblage suggesting that FR‟s system has most likely 575

evolved by closed-system fractional crystallization. At a temperature of 880°C, the magmatic 576

system is subject to a dramatic variation in its physical properties (viscosity, density and water 577

dissolution) as biotite, plagioclase and alkali feldspars start to crystallize. At this temperature, an 578

abrupt decrease in the fraction of melt from 0.5 to 0.1 occurs. The sudden decrease of viscosity and 579

density at this pseudo invariant point temperature and the dramatic change in volume fraction of 580

water from 0.1 to 0.6 is, we speculate, the „trigger‟ mechanism for the eruption of FR magma. 581

Mi1‟s petrological evolution has been simulated by isobaric fractional crystallization. The 582

starting parental composition based on MI‟s in olivine suggests a more primitive parent. The 583

system, containing 2 wt% H2O, has evolved from pressure of 0.3 GPa and oxygen fugacity values 584

around QFM+1. The crystallization sequence is represented by olivine (Tliq = 1300°C), followed by 585

clinopyroxene (T= 1160°C), spinel (T=1070°C), apatite (T=1020°C), plagioclase (T=990°C), 586

biotite (T=960°C) and alkali feldspar (T = 900°C). In the case of Mi1, simulations have not shown 587

invariant temperature behaviour but only a variation of melt fraction (fm) from 0.5 to 0.1 in a 588

temperature span of 90°C (around 990°C), due to the crystallization of alkali feldspars, plagioclase 589

and biotite. A good agreement between observed and calculated mineral compositions suggests that 590

also Mi1 has undergone to a fractional crystallization process even though MIs within later formed 591

clinopyroxene phenocrysts do not appear to represent equilibrium liquids trapped along the liquid 592

25

line of descent. This suggests that reaction between trapped melt and clinopyroxene may be 593

important. The different eruptive style of FR and Mi1 may be related to their different volatile 594

contents in agreement with H2O contents measured by EMPA and SIMS for both eruptions, 595

different melt fraction vs T relationship and eruptive vent location. FR‟s explosive eruption 596

occurred at centre of the CF caldera, while the more “effusive-like” Mi1 eruption occurred along a 597

fissure fracture influenced by the regional fault system in the northern portion of the CF caldera. 598

A general model to explain the mechanism of CF‟s intermediate to small volume pyroclastic 599

eruptions is therefore defined. Magma bodies form in the shallow crust, where they reside for some 600

time, cool and crystallize. Near the end of the crystallization process, water saturation is achieved; 601

the density stratified magma body develops a region (cap) where the volume fraction of fluids 602

exceeds 50% vol. This buoyant cap of water saturated melt (and some dispersed 603

crystals/phenocrysts) coupled with supercritical fluid is mechanically unstable, due to the denser 604

overlying crust (composed by pyrometamorphic rocks, hydrothermally altered sediments, etc) and 605

produces roof hydrofracture and the propagation of volatile-saturated magma–filled cracks. Once 606

the release of pressure during decompression occurs, further exsolution of fluids takes place, along 607

with expansion, shock waves and eruption column formation. 608

The timescale of evolution of FR magmatic system can be constrained based on rates of heat loss 609

from the nearby geothermal system of Mofete, the volume of the system and the difference between 610

the enthalpy at the liquidus and the enthalpy at the lowest melt fraction (fm = 0.05). The results 611

show that FR‟s system has evolved over a time interval of 6.5 ± 3.5 ka, meaning that the magma 612

probably evolved from a basaltic trachy andesitic melt to a trachytic composition over about 6000 613

years. Thermal timescale calculation for Mi1 gives estimate of potential evolving of the system 614

from a basaltic to a trachy andesitic composition in a time span of 2.5 ± 1.5 ka. 615

Similarly with the bradyseism model of Lima et al. (2009), the temporal scale of this quasi 616

periodic behavior is linked with the cooling and crystallization rate of a melt dominated parental 617

magma and can be explained by the periodicity of both volcanic eruptions and bradyseismic events. 618

26

The greater is the mass that cools (i.e. the emplaced magma body) the longer the periods of time to 619

build and trigger the eruptions. Big eruptions such as the Neapolitan Yellow Tuff have required 620

100‟s ka for the magmatic system to reach its dynamic instability, while the time interval of the 621

system development for small eruptions like FR and Mi1 is of order 1-10 ka. At yet shorter time 622

scales of years to decades to centuries, bradyseism is embedded, related to episodes of permeability 623

changes brought on by hydrothermal/thermophysical effects in the wallrock (Bodnar et al., 2007, 624

Lima et al., 2009). 625

CF‟s area represents therefore a unique place to study the effects of magmatic-hydrothermal 626

quasiperiodic behaviour in natural systems characterized by phreatomagmatism, bradyseism and 627

alkaline magma petrogenesis. 628

629

8.1 References 630

Controlla consistenza delle ref con più autori, a volte c’è: and, altre: & altre: la virgola…. 631

Dopo il cognome a volte c’è la virgola a volte no… 632

L’anno di pubblic a volte è iportato fra parentesi a volte no.. 633

insomma devi controllare che le ref siano come le chiede la rivista… 634

635

- Agip, 1987. Modello geotermico de1 sistema flegreo (Sintesi). Servizi Centrali per 636

1‟Esplorazione. SERG-MESG, San Donato, 23 pp. 637

- Anderson, A.T., 1974. Chlorine, sulfur and water in magmas and oceans. Geol. Soc. Amer. Bull. 638

85: 1485–1492. 639

- Anderson, A.T., 1976. Magma mixing: Petrological process and volcanological tool. J. Volcanol. 640

Geotherm. Res. 1: 3-33. 641

- Anderson, A.T., 2003. An introduction to melt (glass±crystals) inclusions. Fluid Inclusions. 642

Analysis and Interpretations. Mineral. Assoc. Canada 32, 374 pp. 643

27

- Anderson, A.T., Davis, A.M., Lu, F., 2000. Evolution of the Bishop Tuff rhyolitic magma based 644

on melt and magnetite inclusions and zoned phenocrysts. J. Petrol. 41: 449-473 645

- Arienzo, I., Civetta, L., Heumann, A., Wörner, G., Orsi, G., 2009. Isotopic evidence for open 646

system processes within the Campanian Ignimbrite (Campi Flegrei–Italy) magma chamber. Bull. 647

Volcanol. 71: 285–300. 648

- Asimow, P.D., Ghiorso, M.S., 1998. Algorithmic modifications extending MELTS to calculate 649

subsolidus phase relations. Am. Mineral. 83: 1127–1131 650

- Barberi, F., Cassano, E., La Torre, P., Sbrana, A., 1991. Structural evolution of Campi Flegrei 651

Caldera in light of volcanological and geophysical data. J. Volcanol. Geotherm. Res. 48: 33-49 652

- Belkin, H.E., De Vivo, B., Roedder, E., Cortini, M., 1985. Fluid inclusions geobarometry from 653

ejected Mt. Somma-Vesuvius nodules. Am. Mineral. 70: 288–303. 654

- Bodnar, R.J., Cannatelli, C., De Vivo, B., Lima, A., Belkin, H.E., Milia, A., 2007. Quantitative 655

model for magma degassing and ground deformation (bradyseism) at Campi Flegrei, Italy: 656

Implications for future eruptions. Geology, 35 (9): 791-794. 657

- Burnham, C.W., 1979. The importance of volatile constituents. In “The evolution of the igneous 658

rocks; 50th

anniversary perspective”. Princeton University Press, 439-482. 659

- Cannatelli, C., Lima, A., Bodnar, R.J., De Vivo, B., Webster, J.D., Fedele, L., 2007. Geochemistry 660

of melt inclusions from the Fondo Riccio and Minopoli 1 eruptions at Campi Flegrei (Italy). Chem. 661

Geol. 237: 418–432 662

- Cashman, K.V., Sturtevant, B., Papale, P., Navon, O., 2000. Magmatic Fragmentation. In: 663

Sigurdsson, H., Houghton, B.F., McNutt, S., Rhymer, H., Stix, J. (Editors) Encyclopedia of 664

Volcanology, Academic Press, San Diego, 419-430. 665

- Carroll, M.R., Holloway, J.R., 1994. Volatiles in magma., Rev. Mineral 30, 517 pp. 666

- Civetta, L., Orsi, G. , Pappalardo, L., Fisher, R.V., Heiken, G., Ort, M., 1997. Geochemical 667

zoning, mingling, eruptive dynamics depositional processes - the Campanian Ignimbrite, Campi 668

Flegrei caldera, Italy. J. Volcanol. Geotherm. Res. 75: 183-219 669

28

- Clocchiatti, R., 1975. Les inclusions vitreuses des cristaux de quartz, Etude optique, 670

hermooptique et chemique, Applications geologiques. Soc. Geol. France, Memoires, New Series 671

122, 96 pp. 672

- Cottrell, E., Spiegelman, M., Langmuir, C. H., 2002. Consequences of diffusive reequilibration for 673

the interpretation of melt inclusions, Geochem. Geophys. Geosyst., 3(5), 674

doi:10.1029/2001GC000205. 675

- Crisp, J.A., 1984. Rates of magma emplacement and volcanic output. J. Volcanol. Geotherm. Res. 676

20: 177-211. 677

- D‟Antonio, M., Civetta, L., Di Girolamo, P., 1999. Mantle source heterogeneity in the Campanian 678

Region (South Italy) as inferred from geochemical and isotopic features of mafic volcanic rocks 679

with shoshonitic affinity. Mineral. Petrol. 67: 163-192 680

- D‟Antonio, M., Tonarini, S., Arienzo, I., Civetta, L., Di Renzo, V., 2007. Components and 681

processes in the magma genesis of the Phlegrean Volcanic District, southern Italy, Spec. Pap. Geol. 682

Soc. Am. 418: 203– 220. 683

- Danyushevsky, L.V., Della Pasqua, F.N., Sokolov, S., 2000. Reequilibration of melt inclusions 684

trapped by magnesian olivine phenocrysts from subduction-related magmas: petrological 685

implications. Contrib. Mineral. Petrol. 138: 68–83. 686

- Deino, A.L., Orsi, G., de Vita, S., Piochi, M., 2004. The age of the Neapolitan Yellow Tuff 687

caldera-forming eruption (Campi Flegrei caldera-Italy) assessed by 40

Ar/39

Ar dating method. J. 688

Volcanol. Geotherm. Res. 133: 157-170 689

- De Astis, G., Pappalardo, L., Piochi, M., 2004. Procida Volcanic History: new insights into the 690

evolution of the Phlegraean Volcanic District (Campania region, Italy). Bull. Volcanol. 66: 622-691

641. 692

- De Lorenzo, S., Gasparini, P., Mongelli, F. & Zollo, A., 2001. Thermal state of the Campi Flegrei 693

Caldera inferred from seismic attenuation tomography. J. Geodyn. 32: 467-486. 694

29

- De Vivo, B., Rolandi, G., Gans, P.B., Calvert, A., Bohrson, W.A., Spera, F.J., Belkin H.E., 2001. 695

New constraints on the pyroclastic eruptive history of the Campanian volcanic Plain (Italy). 696

Mineral. Petrol. 73: 121-143. 697

- De Vivo, B., Bodnar ,R.J., 2003. Melt inclusions in volcanic systems. Methods, applications and 698

problems. Series Developments in Volcanology 5, Elsevier, Amsterdam, pp. 250. 699

- De Vivo, B., Lima A., Webster, J. D., 2005. Volatiles in magmatic-volcanic systems. Elements 1: 700

19-24. 701

- De Vivo, B., Lima, A., 2006. A hydrothermal model for ground movements (bradyseism) at 702

Campi Flegrei, Italy. In: De Vivo (Editor), Volcanism in the Campania Plain: Vesuvius, Campi 703

Flegrei and Ignimbrites. Developments in Volcanology 9, Elsevier, 289-317. 704

- Di Girolamo, P., Ghiara, M.R., Lirer, L., Munno, R., Rolandi, G., Stanzione, D., 1984. 705

Vulcanologia e Petrologia dei Campi Flegrei. Boll. Soc. Geol. It. 103: 349-413 706

- Di Vito, M.A., Lirer, L., Mastrolorenzo, G., Rolandi, G., 1987. The 1538 Monte Nuovo eruption 707

(Campi Flegrei, Italy). Bull. Volcanol. 49: 608-615. 708

- Di Vito, M.A., Isaia, R., Orsi, G., Southon, J., de Vita, S., D‟Antonio, M., Pappalardo, L., Piochi, 709

M. 1999. Volcanism and deformation since 12,000 years at the Campi Flegrei caldera (Italy). J. 710

Volcanol. Geotherm. Res. 91: 221-246. 711

- Di Vito, M.A., Sulpizio, R., Zanchetta, G., D'Orazio, M., 2008. The late Pleistocene pyroclastic 712

deposits of the Campanian Plain: new insights into the explosive activity of Neapolitan volcanoes. 713

J. Volcanol. Geotherm. Res. 17: 19–48 714

- Fowler, S.J., Spera, F.J., Bohrson, W.A., Belkin, H.E., De Vivo, B. 2007. Phase equilibria 715

constraints on the chemical and physical evolution of the Campanian Ignimbrite. J. Petrol. 48 (3): 716

459-493 717

- Fowler, S.J., Spera, F.J., 2008. Phase equilibria trigger for explosive volcanic eruptions, Geophys. 718

Res. Lett. 35, L08309, doi:10.1029/2008GL033665. 719

30

- Frezzotti, M.L., 2001. Silicate–melt inclusions in magmatic rocks: application to petrology. Lithos 720

55: 273–299. 721

- Gaetani, G. A., Watson, E. B., 2000. Open system behavior of olivine-hosted melt inclusions, 722

Earth Planet. Sci. Lett. 183: 27–41. 723

- Ghiorso, M.S., Sack, R.O. 1995. Chemical mass transfer in magmatic processes IV. A revised and 724

internally consistent thermodynamic model for the interpolation and extrapolation of liquid–solid 725

equilibria in magmatic systems at elevated temperatures and pressures. Contrib. Mineral. Petrol. 726

119:197–212. 727

- Kress, V. C., Ghiorso, M. S. 2004. Thermodynamic modeling of post-entrapment crystallization in 728

igneous phases. J. Volcanol. Geotherm. Res. 137: 247-260. 729

- Kohlstedt, D.L.,Holtzman, B.K., 2009. Shearing melt out of the Earth; an experimentalist‟s 730

perspective on the influence of the deformation on melt extraction. Annu. Rev. Earth Planet. Sci. 731

37:561-593. 732

- Lange, R. A., 1994. The effect of H2O and CO2 on the density and viscosity of silicate melts. In: 733

Volatiles in Magmas. Rev. Mineral. 30, pp 331-370. 734

- Lowenstern, J.B., 1994. Dissolved volatile concentrations in ore-forming magma. Geology 22: 735

893–896. 736

- Lima, A., De Vivo, B., Spera, F.J., Bodnar,R.J., Milia, A., Nunziata, C., Belkin, H.E., Cannatelli, 737

C., 2009. Thermodynamic model for uplift and deflation episodes (bradyseism) associated with 738

magmatic-hydrothermal activity at the Campi Flegrei (Italy). Earth Sci.Rev. 97: 44-58. 739

- Marianelli, P., Sbrana, A., Proto, M., 2006. Magma chamber of the Campi Flegrei supervolcano at 740

the time of eruption of the Campanian Ignimbrite. Geology 34 (11): 937–940. 741

- Metrich, N., Wallace, P.J., 2008. Volatiles abundance in basaltic magmas and their degassing 742

paths tracked by melt inclusions. Rev. Mineral. Geochem. 69: 363-402 743

31

- Michael, P.J., McDonough, W.F., Nielsen, R.L., Cornell, W.C., 2002. Depleted melt inclusions in 744

MORB plagioclase: messages from the mantle or mirages from the magma chamber? Chem. Geol. 745

183: 43–61. 746

- Moore, G., 2008. Interpreting H2O and CO2 contents in melt inclusions: constraints from solubility 747

experiments and modelling. In: K.D. Putirka, F.J. Tepley III (Editors) Minerals, Inclusions and 748

Volcanic Processes, Rev. Mineral. 69: 333–361 749

- Moore, G., Carmichael, I.S.E. 1998. The hydrous phase equilibria (to 3kbar) of an andesite and 750

basaltic andesite from western Mexico: constraints on water content and conditions of phenocrysts 751

growth. Contrib. Mineral. Petrol. 130: 304-319 752

- Morgan, G.B., London, D., 1996. Optimizing the electron microprobe analysis of hydrous alkali 753

aluminosilicate glasses. Am. Mineral. 81: 1176–1185. 754

- Ochs, F.A., Lange, R.A., 1999. The density of hydrous magmatic liquids. Science 283: 1314-1317 755

- Orsi, G., de Vita, S., Di Vito, M., 1996. The restless, resurgent Campi Flegrei nested caldera 756

(Italy): constraints on its evolution and configuration. J. Volcanol. Geotherm. Res. 74: 179-214 757

- Pappalardo, L., Civetta, L., D‟Antonio, M., Deino, A.L., Di Vito, M.A., Orsi, G., Carandente, A., 758

de Vita, S., Isaia, R., Piochi, M., 1999. Chemical and isotopical evolution of the Phlegraean 759

magmatic system before the Campanian Ignimbrite (37 ka) and the Neapolitan Yellow Tuff (12 ka) 760

eruptions. J. Volcanol. Geotherm. Res. 91: 141-166 761

- Pappalardo, L., Piochi, M., D‟Antonio, M., Civetta, L., Petrini, R., 2002. Evidence for multi-stage 762

magmatic evolution during the past 60 ka at Campi Flegrei (Italy) deduced from Sr, Nd and Pb 763

isotope data. J. Petrol. 43: 1415-1434 764

- Peccerillo, A. 2005. Plio-Quaternary Volcanism in Italy. Springer, 365 pp. 765

- Pichou, J.L., Pouchoir, F., 1985. “PAP” procedure for improved quantitative microanalysis. 766

Microbeam Anal. 20, 104-105. 767

- Roach, A.L., 2005. The evolution of silicic magmatism in the post-caldera volcanism of the 768

Phlegrean Fields, Italy. Ph.D. thesis, Brown University, Providence, RI, 151 pp. 769

32

- Roedder, E., 1979. Origin and significance of magmatic inclusions. Bull. Mineral. 102: 487–510. 770

- Roedder, E., 1984. Fluid inclusions. Rev. Mineral. 12, 644 pp. 771

-Roggensack, K.. 2001. Sizing up crystals and their melt inclusions: a new approach to 772

crystallization studies. Earth Planet. Sci. Lett. 187: 221-237. 773

- Rolandi, G., Bellucci, F., Heizler, M.T., Belkin, H.E., De Vivo, B., 2003. Tectonic controls on the 774

genesis of ignimbrites from the Campanian Volcanic Zone, southern Italy. Mineral. Petrol, 79: 3 - 775

31. 776

- Rollinson, H.R., 1993. Using Geochemical Data: evaluation, presentation, interpretation. 777

Longman/Wyllie. Harlow/New York, 352 pp. 778

- Rosi, M., Sbrana, A., 1987, Phlegraean Fields. Quaderni de “La Ricerca Scientifica”: Consiglio 779

Nazionale delle Ricerche Monograph 114, Volume 9, 175 pp. 780

- Shimizu, N., Hart, S.R., 1982. Application of the ion probe to geochemistry and cosmochemistry. 781

Annu. Rev. Earth Planet. Sci. 10: 483–526. 782

- Sobolev, A.V., 1996. Melt inclusions in minerals as a source of principle petrological information. 783

Petrol. 4: 209–220. 784

- Sobolev, A.V., Kamenetsky, V.S., Metrich, N., Clocchiatti, R., Koronova, N.N., Devirts, A.L., 785

Ustinov, V.I., 1990. Volatile regime and crystallization conditions in Etna hawaiite lavas. Geochem. 786

Int.: 53–65. 787

- Sparks, R.S.J., Barclay, J., Jaupart, C., Mader, H.M., Phillips, J.C. 1994. Physical aspects of 788

magmatic degassing I. Experimental and theoretical constraints on vesciculation. Rev. Mineral. 30: 789

413-446 790

- Spera, F. J., Stein, D. J., Lejeune, A. M., Bottinga, Y., Trull, T. W., Richet, P., 2000. Rheology of 791

bubble-bearing magmas; discussion and reply [modified]. Earth Planet. Sci. Lett. 175: 327-334. 792

- Tonarini, S., Leeman, W.P., Civetta, L., D'Antonio, M., Ferrara, G., Necco, A., 2004. B/Nb and 793

δ11

B systematics in the Phlegrean Volcanic District (PVD). J. Volcanol. Geotherm. Res. 133: 123–794

139. 795

33

- Tonarini, S., D'Antonio, M., Di Vito, M.A., Orsi, G., Carandente, A. 2009. Geochemical and B–796

Sr–Nd isotopic evidence for mingling and mixing processes in the magmatic system that fed the 797

Astroni volcano (4.1–3.8 ka) within the Campi Flegrei caldera (southern Italy). Lithos 107: 135-798

151. 799

- Wallace, P.J., 2005. Volatiles in subduction zone magmas; concentrations and fluxes based on 800

melt inclusion and volcanic gas data. J. Volcanol. Geotherm. Res. 140: 217-240 801

- Watson, E. B. 1976. Glass inclusions as samples of early magmatic liquid; determinative method 802

and application to South Atlantic basalt. J. Volcanol. Geotherm. Res. 1: 73-84. 803

- Webster, J.D., Burt, D.M., Aguillon, R.A., 1996. Volatile and lithophile trace-element 804

geochemistry of heterogeneous Mexican tin rhyolite magmas deduced from compositions of melt 805

inclusions. Geochim. Cosmochim. Acta 60: 3267–3283. 806

- Webster, J.D., Raia, F., De Vivo, B., Rolandi, G., 2001. The behavior of chlorine and sulfur 807

during differentiation of the Mt. Somma– Vesuvius magmatic system. Mineral. Petrol. 73: 177–201. 808

- White, S.M., J.A. Crisp, F.J. Spera, 2006. Long-term volumetric eruption rates and magma 809

budgets, Geochem. Geophys. Geosyst. 7, Q03010, doi:10.1029/2005GC001002. 810

- Wohletz, K., Civetta, L. Orsi, G. (1999). Thermal evolution of the Phlegraean magmatic system. J. 811

Volcanol. Geotherm. Res. 91: 381-414. 812

- Zollo, A., Maercklin, N., Vassallo, M., Dello Iacono, D., Virieux, J., Gasparini P., 2008. Seismic 813

reflections reveal a massive melt layer feeding Campi Flegrei caldera, Geophys. Res. Lett. 35, 814

L12306, doi:10.1029/2008GL034242. 815

816

817

818

819

820

821

822

823

824

825

34

Figure Captions 826

Fig. 1 Schematic geological map of Campi Flegrei Volcanic District (CFVD), FR = Fondo Riccio, 827

Mi1= Minopoli 1. Modified after Orsi et al. (1996) and Peccerillo (2005). 828

Fig. 2 CaO-MgO-Al2O3 triangular diagrams (Watson, 1976) for melt inclusions hosted in olivine 829

(circles) and pyroxene (squares). Point “I” is the intersection of olivine and pyroxene fractionation 830

lines and represent the composition of the magmatic liquid at the time of melt inclusion formation. 831

(a) Fondo Riccio, (b) Minopoli 1. 832

Fig. 3 Total Alkali Silica diagram (Le Bas et al., 1986) showing the MELTS results simulations for 833

variable fO2. Symbols are shown in the legend. Shaded area represents field for all Fondo Riccio 834

melt inclusions data (Cannatelli et al., 2007). 835

Fig.4 Phase proportion as a function of temperature for MELTS simulation at variable oxygen 836

fugacity, P = 0.2 GPa and H2O = 2wt% for Fondo Riccio. Ap = apatite, Bio = biotite, Cor = 837

corundum, Cpx =clinopyroxene, Ksp = alkali feldspar, Leu = leucite, Ol = olivine, Plag = 838

plagioclase feldspar, Rut = rutile, Sp = spinel. (a) QFM+2, (b) QFM+1, (c) QFM, (d) QFM-1, (e) 839

QFM-2. 840

Fig. 5 Minopoli 1 calculated compositions for fO2 = QFM+1, H2O = 2wt% and P = 0.2 GPa and 0.3 841

GPa. (a) Feldspar, (b) Spinel. 842

Fig. 6 Fondo Riccio‟s phase proportion diagrams as a function of temperature for MELTS 843

simulation at variable oxygen fugacity, P = 0.15 GPa and H2O = 3wt%. Ap = apatite, Bio = biotite, 844

Cpx =clinopyroxene, Ksp = alkali feldspar, Ol = olivine, Plag = plagioclase feldspar, Rh-ox = 845

rhombohedral oxide, Sp = spinel. (a) QFM+1, (b) QFM. 846

Fig. 7 Oxides diagram for Fondo Riccio in the best cases of 0.15 GPa, 3wt% H2O and varying fO2 847

between QFM and QFM+1. (a) FeOtot, (b) Al2O3, (c) Na2O, (d) K2O, (e) SiO2 and (f) CaO . 848

Fig. 8 Fondo Riccio‟s calculated mineral compositions for the best case of MELTS simulations. (a) 849

Circles = olivine, diamonds = clinopyroxene; (b) triangles = feldspar; (c) triangles = spinel. Grey 850

symbols are MELTS generated data, open symbols are mineral data collected by EMPA. 851

35

Fig. 9 Minopoli 1 phase proportion as a function of temperature for MELTS simulation at variable 852

oxygen fugacity, P = 0.3 GPa and H2O = 2wt%. Ap = apatite, Bio = biotite, Cor = corundum, Cpx 853

=clinopyroxene, Ksp = alkali feldspar, Leu = leucite, Ol = olivine, Plag = plagioclase feldspar, Rut 854

= rutile, Sp = spinel. (a) QFM, (b) QFM+1. 855

Fig. 10 Oxides diagram for Minopoli 1 in the best case produced by MELTS, P = 0.3 GPa, 2wt% 856

H2O and fO2 = QFM. (a) SiO2, (b) CaO, (c) FeOtot

, (d) Al2O3, (e) Na2O and (f) K2O. 857

Fig. 11 Minopoli 1‟s calculated mineral compositions for the best case of MELTS simulations. (a) 858

Circles = olivine, diamonds = cpx; (b) triangles = feldspar. Grey symbols are MELTS generated 859

data, open symbols are mineral data collected by EMPA. 860

Fig. 12 Variation of melt physical properties for Fondo Riccio along the liquid line of descent for 861

the case P= 0.15 GPa, 3wt% H2O and QFM+1. (a) Density of melt versus T, (b) Dissolved water 862

content versus T, (c) melt viscosity versus T, (d) volume fraction of water versus T. H2O saturates 863

at 1070°C. 864

Fig, 13 Variation of melt physical properties for Minopoli 1 along the liquid line of descent for the 865

case P= 0.3 GPa, 2wt% H2O and QFM. (a) Density of melt versus T, (b) Dissolved water content 866

versus T, (c) melt viscosity versus T, (d) volume fraction of water versus T. there is no saturation of 867

water along the liquid line of descent for this case. 868

Fig. 14 Timescale evolution and temporal crystallization history showed by phase proportion in 869

function of magma temperature. (a) Fondo Riccio, (b) Minopoli 1. According to the best fit model, 870

ages represent time before each eruption when specific phase mineral begin crystallizing. 871