Carbonate-derived CO2 purging magma at depth: Influence on the eruptive activity of Somma-Vesuvius, Italy

Earth and Planetary Science Letters 310 (2011) 84–95

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Earth and Planetary Science Letters j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / e p s l

Carbonate-derived CO2 purging magma at depth: Influence on the eruptive activity of Somma-Vesuvius, Italy Luigi Dallai a,⁎, Raffaello Cioni b, c, Chiara Boschi a, Claudia D'Oriano c a b c

CNR-Istituto di Geoscienze e Georisorse, Via Moruzzi 1, 56124 Pisa, Italy Dip. to Scienze della Terra, Via Trentino 51, 09127 Cagliari, Italy INGV, sezione di Pisa, Via della Faggiola 32, 56126 Pisa, Italy

a r t i c l e

i n f o

Article history: Received 31 January 2011 Received in revised form 13 July 2011 Accepted 14 July 2011 Available online 17 September 2011 Editor: R.W. Carlson Keywords: stable-isotope magma geochemistry CO2-degassing Vesuvius

a b s t r a c t Mafic phenocrysts from selected products of the last 4 ka volcanic activity at Mt. Vesuvius were investigated for their chemical and O-isotope composition, as a proxy for primary magmas feeding the system. 18O/16O ratios of studied Mg-rich olivines suggest that near-primary shoshonitic to tephritic melts experienced a flux of sedimentary carbonate-derived CO2, representing the early process of magma contamination in the roots of the volcanic structure. Bulk carbonate assimilation (physical digestion) mainly occurred in the shallow crust, strongly influencing magma chamber evolution. On a petrological and geochemical basis the effects of bulk sedimentary carbonate digestion on the chemical composition of the near-primary melts are resolved from those of carbonate-released CO2 fluxed into magma. An important outcome of this process lies in the effect of external CO2 in changing the overall volatile solubility of the magma, enhancing the ability of Vesuvius mafic magmas to rapidly rise and explosively erupt at the surface. © 2011 Elsevier B.V. All rights reserved.

1. Introduction Significant interaction between mafic magma and crust has been documented at volcanic systems set in carbonate basements (e.g. Vesuvius, Italy; Popocatepetl, Mexico; Merapi, Indonesia; Ayuso et al., 1998; Goff et al., 2001; Chadwick et al., 2007). Carbonate assimilation occurs to variable extents during magma crystallization and evolution at shallow crustal levels, modifying the overall composition of the residual melts (Barnes et al., 2005; Dallai et al., 2004; Freda et al., 2008; Iacono Marziano et al., 2007; 2008; Mollo et al., 2010). Much less is known about processes controlling the interaction between mantle-derived melts and carbonate, and about the effects of carbonate-derived CO2 on primitive magmas during their ascent to the surface (Deegan et al., 2010). Understanding the mechanisms of such carbonate/melt interaction is important to resolve the effects of carbonate digestion from those of carbonate-released CO2 fluxing into magma, and their possible influence on the eruptive style. Thermal decomposition of carbonate (Stanmore and Gillot, 2005) and metamorphic–metasomatic reactions (Nabeleck, 2007) represent the “type-mechanisms” of magma–carbonate interaction (Baker and Black, 1980; Gaeta et al., 2009; Wenzel et al., 2002), during which a massive exchange of heat and mass occurs, producing large amounts of CO2 available for direct fluxing through the magma. While continu⁎ Corresponding author. Tel.: + 39 0503152315; fax: + 39 0503152323. E-mail addresses: [email protected] (L. Dallai), [email protected] (R. Cioni), [email protected] (C. Boschi), [email protected] (C. D'Oriano). 0012-821X/$ – see front matter © 2011 Elsevier B.V. All rights reserved. doi:10.1016/j.epsl.2011.07.013

ously fluxing the magma at depth, CO2 has the potential to change the composition and solubility of volatile components dissolved in the melt (Dixon and Stolper, 1995; Papale, 1999), promoting an increase of magma explosivity and introducing substantial variations in the expected eruption scenarios defined for hazard assessment. The influence of magma–carbonate interaction on the composition of volcanic products from Mt. Somma-Vesuvius (SV), one of the most hazardous volcanoes on a world-wide scale, has been repeatedly debated in the last decades (e.g. Ayuso et al., 1998; Piochi et al., 2006; Rittmann, 1933; Savelli, 1967). Among these Authors, Rittmann (1933) first proposed the idea that the peculiar composition of the rocks was related to the effects of carbonate assimilation in the Mesozoic basement of the volcano. The idea was successively discarded by Savelli (1967), and recently reviewed and proposed on the base of isotopic data (Ayuso et al., 1998; Civetta et al., 2004; Di Renzo et al., 2007; Piochi et al., 2006) and of experimental petrology data (Iacono Marziano et al., 2007, 2008; Freda et al., 2008; Gaeta et al., 2009; Mollo et al., 2010). The occurrence of skarn and metasomatized cumulatic ejecta in the products of the major SV explosive eruptions (Barberi and Leoni, 1980), and the presence of exotic, silica-poor, alkali and Ca-rich melt inclusions in mafic crystals of some recent eruptions (Fulignati et al., 2001), unequivocally indicate that interaction between SV magmas and carbonate rocks was pervasive during magma residence and crystallization within shallow reservoirs (less than 8–10 km; Auger et al., 2001; De Natale et al., 2006; Scaillet et al., 2008). The effects of carbonate interaction on mafic magmas feeding the shallow system are instead less defined,

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mainly due to the general absence of primary (or nearly primary) magmas among the SV eruptive products. In recent years, these effects have been investigated using experimental petrology (Iacono Marziano et al., 2008) suggesting that alkali-rich, tephritic magmas may be derived from primary shoshonitic melts by assimilation of 10 to 20 wt.% of carbonate. Due to the large difference in 18O/ 16O ratios between sedimentary carbonate (formed in a low-temperature shallow environment) and mafic (mantle-derived) magmas, oxygen isotope systematics are a powerful tool for tracing the process of carbonate addition to nearprimary K-rich melts and pyroclastic materials (Dallai et al., 2004; Frezzotti et al., 2007; Gaeta et al., 2006). In particular, rapid cooling rate of pyroclasts prevents subsolidus re-equilibration of O-isotope composition, and pre-eruptive compositions are likely preserved (e.g. Eiler et al., 1997). The present work focuses on the oxygen isotope composition of selected mafic crystals extracted from pyroclastic products of the last 4 ka of SV activity, in order to constrain the role and processes of magma–carbonate interaction in producing the compositional variability observed in the “primary” magmas (from shoshonites to tephrites), and possibly affecting their explosivity. 1.1. The Somma-Vesuvius magmatic activity Magmatism in the SV area, part of the potassic Quaternary Campania Province (Southern Italy), has been generally interpreted as related to magma generation in an upper mantle contaminated by material coming from the West-directed subducting Adria-Ionian plates (Peccerillo and Lustrino, 2005). An alternative hypothesis invokes a mantle plume rising beneath the southern Tyrrhenian Sea and contamination by subduction (Gasperini et al., 2002). The early magmatic products in the SV area date back at about 400 ka (Brocchini et al., 2001). A mainly effusive volcanic activity occurred discontinuously till about 20 ka, leading to the formation of Mt. Somma stratocone. After this period, the magmatic activity changed to prevalently explosive, and in the period up to 79 AD, four caldera-forming Plinian eruptions, and several subplinian to mid-intensity ash-dominated events occurred (Cioni et al., 2008). The activity of the last 2 ka was characterized by periods of low to midintensity, mafic explosive eruptions, alternated with periods of mainly effusive activity. In this period, the summit cone of Vesuvius was erected. During the last 20 ka, magma composition changed progressively shifting from nearly saturated alkaline melts towards more alkalirich silica undersaturated products (e.g. Ayuso et al., 1998; Cioni et al., 2008; Santacroce et al., 2008). Effusive and explosive products are represented by moderately to highly evolved compositions, mainly varying from trachybasalts to trachytes, from phono-tephrites to phonolites, and from tephrite/basanite to foidites. The substantial absence of primitive products is characteristic of SV products, and direct information on primary mantle-derived melts can be only derived from the study of silicatic melt inclusions in mafic minerals (forsteritic olivine and diopsidic pyroxene) occurring as xenocrysts in many erupted products (Cioni et al., 1998; Marianelli et al., 1995; Marianelli et al., 2005). These xenocrysts have been interpreted as derived from the crystallization of primitive magmas during their ascent or as physical mixing with the most evolved, cooler magmas residing in shallower reservoirs. Using the volatile (H2O and CO2) content of these melt inclusions as a proxy for the pressure of crystallization of the hosting minerals, pressure of about 200–300 MPa is derived (Marianelli et al., 2005). Geophysical and experimental petrology data are in agreement with these findings, suggesting the presence of a large volume reservoir at about 8–10 km depth (Auger et al., 2001; Scaillet et al., 2008). The whole volcanic sequence rests on a pile of tectonic units dominated by a very thick (possibly more than 10 km, according to Patacca and Scandone, 2007) deposit of Mesozoic carbonates (limestones and dolomitic limestones) arranged in a duplex system. Data from a geothermal borehole drilled on the southern slope of the volcano highlight that the first SV magmatic products were inter-

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layered with a neritic to continental Pleistocene silico-clastic sequence, which directly covers the Mesozoic carbonatic basement at a depth of about 1900 m (Brocchini et al., 2001). Data on metamorphic and non-metamorphic carbonate ejecta from the Plinian eruptions (Barberi and Leoni, 1980) suggest the predominance of limestones and dolomitic limestones over pure dolomites. 2. Samples Magmatic mafic xenocrysts of SV products, were investigated for their oxygen isotope and major and trace elements compositions. Single high-Mg olivine and diopside crystals from the Plinian products of the tephri-phonolite to phonolite Avellino (3.9 ka BP; Sulpizio et al., 2010) and Pompeii Pumice (AD 79; Cioni et al., 1995) eruptions, from the subplinian, phono-tephrite to tephri-phonolite, Pollena (AD 472; Sulpizio et al., 2005) eruption, and from the phono-tephrite products of a violent strombolian eruption which occurred in the 8th Century from a lateral vent in the south-western sector of the volcano (AS2f; Cioni et al., 2008). Despite the large different compositions shown by these 4 eruptions, diopside and forsteritic olivine are ubiquitous in their products, although in different proportions, pointing out the important role of mixing processes during magma residence in a shallow reservoir, and of magma extraction during eruption (Cioni et al., 1995; Sigurdsson et al., 1990). Diopside and forsteritic olivine have been also described at SV as fundamental mineral phases in skarns and thermometamorphic rocks (Barberi and Leoni, 1980). However, “non-magmatic”, Mg-rich, olivine and clinopyroxene have major and trace elements compositions significantly different from those of “magmatic” phases (Gilg et al., 2001; Fig. 1a). Chemical composition was used here in order to characterize the crystals selected for isotope analyses. In fact, only large euhedral crystals with a clearly magmatic derivation, bordered by glass rims were chosen for the study. Finally, crystals with a large amount of melt inclusions were discarded, in order to avoid any influence on the isotopic composition. 3. Analytical methods Crystals were accurately separated from selected juvenile material of the studied eruptions. Pumice and large scoriae clasts were crushed in a steel mortar and then sieved. Crystals of olivine and diopside between 2 and 1 mm and 1 and 0.5 mm were hand picked under the stereomicroscope and glued with a thermoplastic cement on a glass slide. Samples showing homogeneous color (unzoned crystals) and without inclusions were selected for the successive investigations from a total number of about 100 crystals collected. Composition was determined at the Institute for Mineralogy and Petrology (ETH-Zurich, Switzerland) using a JEOL JXA-8200 electron microprobe with operating conditions of 15 kV accelerating potential, 20 nA current, and 1–10 μm beam size. Estimated precision ranges around 0.05 wt.%. The detection limit is better than 0.01 wt.% for each element. Major elements compositions of olivine and cpx of the AS2f eruption were analyzed by energy dispersive X-ray with an EDAX X4I, using a Philips XL30 scanning electron microscope at the University of Pisa, at an accelerating voltage of 20 kV, beam current of 0.1 nA and working distance of 10 mm. At least 2 analyses per crystal were carried out in order to avoid using strongly zoned crystals. On the same crystals trace elements analyses were performed at the IGG-CNR laboratory of Pavia, using a laser system consisting of a Brilliant, Quantel, Q-switched Nd:YAG laser, working at a wavelength of 266 nm (Tiepolo et al., 2003). The ablated material was carried by an argon–helium mix to a Perkin-Elmer DRC-e ICP-MS. The laser was operated at a repetition rate of 10 Hz, a power of 3.0 mW and spot size of 40 μm. Masses were acquired in peak hopping mode with a dwell time of 10 ms. Nist SRM 610 and 43Ca were adopted as external and internal standards, respectively. Precision and accuracy were evaluated on the USGS-BCR-2 reference material and are estimated to be

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magma mixing (Cioni et al., 1998). In order to study crystals representative of the first phases of magma crystallization, we also selected unzoned, diopside crystals (En45–48 Fs4–9; Cioni et al., 1998). Melt inclusions in diopside are in fact indicative of a very early crystallization, both for their composition and for the high homogeneization temperature measured (Cioni, 2000; Cioni et al., 1998; Marianelli et al., 1995). All the analyzed crystals are similarly LREEenriched (Table 2; Fig. 1) and show convex-upward REE patterns, typical of diopside crystallized from mafic, primitive magma (e.g. Dioh et al., 2009; Ying et al., 2006). The magmatic derivation of diopside is also confirmed by their content in Wollastonite molecule, lower than that typical of skarn material (Gilg et al., 2001; Fig. 1a). An important issue of this study is related to the assumption that the selected olivine–pyroxene pairs were in chemical equilibrium. Melt inclusions hosted in olivine and diopside show similar, homogeneous compositions, thereby implying that crystals are formed in the same kind of melt (Marianelli et al., 1995). Moreover, Fe–Mg partition between olivine and diopside (Kdol/cpxFe/Mg) suggest a crystallization temperature around 1200 °C (based on the geothermometric relationship proposed by Loucks (1996)), in good agreement with the homogenization temperature measured for melt inclusions (varying between 1160 and 1200 °C; Cioni et al., 1998; Marianelli et al., 1995). 4.2. Oxygen isotope composition

Fig. 1. a) Plot of clinopyroxene major element compositions within a portion of the Ca–Mg–Fe triangle. White boxes represent the analyses performed in this work. For comparison we reported also the compositions of Vesuvian magmatic clinopyroxenes analyzed in literature (gray field: Cioni et al., 1998; Landi et al., 1999; Marianelli et al., 1999; Cioni, 2000; Morgan et al., 2004); and skarn clinopyroxene compositions (circles: Gilg et al., 2001). b) REE patterns of diopside clinopyroxene from Mt. Vesuvius eruptions normalized to chondrite composition (Boynton, 1984).

better than 5 and 10% relative, respectively; 2 to 3 analyses were performed for each crystal and the average values were considered. Oxygen isotope compositions of single mineral grains were measured at the CNR-IGG, Pisa, by conventional laser fluorination (Sharp, 1995), reacting the samples under an F2 gas atmosphere. Purified oxygen gas was directly transferred into a Thermo Finnigan Delta XP Isotope Ratio Mass Spectrometer via a 13A zeolite molecular sieve. All the data are given following the standard δ-notation relative to SMOW (Standard Mean Oceanic Water). Duplicate measurements were performed when sufficient material was available and the average δ 18O values were considered ± the standard error of the mean. In the course of analysis, an in-house laboratory QMS quartz standard (δ 18O SMOW = 14.05‰) calibrated vs. the international quartz standard NBS28 (δ 18O = + 9.58‰) was used, yielding an average δ 18O value = + 14.08‰, (1 s = 0.14, n = 12). Standard NBS30 (δ 18O = + 5.24‰) was also used during the study and gave an average value of δ 18O = 5.22‰ (1 s = 0.16; n = 7). 4. Results 4.1. Mineral chemistry Olivine crystals from all the selected eruptions vary in a restricted range of composition (from Fo86 to Fo91), with the most primitive crystals (Fo91) measured in samples collected from the AS2f eruption (Table 1). Olivine has very low trace elements concentrations except for highly compatible elements like Cr and Ni, as expected for crystals growing from basaltic melts. Pyroxene composition at SV is largely variable, reflecting the combined effect of complex processes of fractional crystallization and

The δ 18O values of the measured olivine from the different eruptions at Mt. Vesuvius vary from 5.5 to 7.1‰ (Table 1, Fig. 2a), and variability is narrower within a single eruption. Olivine crystals from Avellino and Pollena eruptions have the largest range (from 5.89 to 7.11‰ and 5.92 to 7.03‰, respectively), while olivine crystals from the Middle Age AS2f eruption vary from 6.04 to 6.52‰, and those from Pompeii eruption from 5.51 to 6.19‰. Overall, no typical mantle δ 18O values (δ 18Ool = 5.18 ± 0.28‰; Mattey et al., 1994) were recovered, and olivine crystals from the Pompeii eruption show a few values similar to the melts of island arc volcanics (Bindeman et al., 2005). This is not unexpected, considering the complex and recent subduction-related volcanic history of Italian Quaternary lavas (e.g. Peccerillo, 1999). Mantle-like O-isotope compositions have been recovered only in a few monomineralic cumulates (Dallai et al., 2004; Peccerillo et al., 2004) suggesting that slight mantle O-isotope variability is overprinted significantly by processes that occurred in the magma chamber(s). Also clinopyroxenes show δ 18O values (Table 2, Fig. 2b) varying over a narrow range (Avellino from 6.25 to 6.77‰; Pollena 6.55 to 6.90‰; AS2f 6.48 to 6.98‰), with crystals from Pompeii eruption showing slightly lower values (6.04 to 6.80‰). The highest δ 18O values (three δ 18O values above 7.1‰) are shown by diopside crystals from the phono-tephritic portion (grey pomice) of the AD 79 Pompeii Pomice eruption. Considering the mean values of each olivine and clinopyroxene population, and their standard deviations (st.dev./sq.root n_samples), the δ 18O values of the largest eruptions (Avellino and Pompeii) do not overlap, whereas those of the smaller eruptions (Pollena and AS2f) vary in the same range (Fig. 3). These data suggest that the near-primary melts from which the minerals crystallized possibly underwent variable contamination during the early stages of crystallization. The fact that the Fo-richest olivine belongs to a small eruption and has high δ 18O values indicates that the almost unevolved mafic magma was modified for its O-isotope composition by interaction with an 18O-enriched phase. The variability measured in the isotopic composition contrasts with the homogeneous major and trace elements composition of the phenocrysts. With the exception of Avellino crystals, the isotopic composition of clinopyroxene and olivine suggests isotopic equilibrium between the two phases (δ 18Odiopside–olivine = 0.4‰; Mattey et al., 1994), and their O-isotope fractionation defines a temperature of 1240 °C (Chiba et al., 1989), slightly higher than the temperature of crystallization based on petrologic inferences (Cioni et al., 1999).

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Table 1 Major elements compositions (wt.%) of olivine in studied eruptions. For each analyzed crystal, Fo (mol) and δ18O are presented. (wt%)

Avellino

SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O Cr2O3 NiO

40.79 0.00 0.01 10.15 0.20 48.14 0.28 0.01 0.00 0.02 0.17

40.76 0.03 0.02 11.13 0.26 46.84 0.25 0.00 0.00 0.01 0.10

40.59 0.06 0.03 12.85 0.22 45.86 0.24 0.00 0.00 0.00 0.14

40.34 0.03 0.02 13.22 0.25 45.78 0.23 0.01 0.00 0.00 0.13

41.30 0.02 0.01 11.06 0.29 46.81 0.28 0.01 0.00 0.00 0.15

40.70 0.00 0.05 10.35 0.23 47.45 0.27 0.03 0.00 0.01 0.15

40.70 0.02 0.01 10.11 0.22 48.02 0.29 0.00 0.00 0.03 0.20

40.23 0.00 0.01 14.91 0.28 45.62 0.20 0.02 0.03 0.02 0.16

40.28 0.01 0.03 13.39 0.23 46.57 0.28 0.01 0.01 0.01 0.16

40.30 0.01 0.02 12.97 0.23 46.86 0.29 0.01 0.00 0.01 0.15

40.55 0.01 0.01 12.76 0.22 46.72 0.29 0.03 0.01 0.02 0.16

40.00 0.01 0.02 12.67 0.22 47.51 0.30 0.02 0.00 0.01 0.15

0.89 5.89 0.10

0.88 6.02

0.87 6.18

0.86 6.17

0.88 7.11

0.89 6.72

0.89 6.49

0.84 6.08 0.06

0.85 5.99

0.86 5.67

0.86 5.75

0.86 5.79

Fo (mol) δ18O (‰) sem

Pompeii

(wt%)

Pompeii

SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O Cr2O3 NiO

40.82 0.01 0.02 10.58 0.20 48.90 0.29 0.03 0.02 0.03 0.20

40.47 0.00 0.02 10.08 0.19 49.51 0.29 0.01 0.01 0.04 0.19

40.55 0.00 0.03 10.33 0.18 49.20 0.30 0.01 0.00 0.03 0.20

40.44 0.00 0.02 10.31 0.18 49.58 0.29 0.01 0.00 0.02 0.20

40.33 0.01 0.02 10.38 0.18 49.12 0.29 0.00 0.01 0.03 0.22

40.95 0.01 0.02 10.31 0.18 48.88 0.30 0.01 0.00 0.02 0.20

40.66 0.01 0.02 12.17 0.20 47.50 0.30 0.03 0.00 0.01 0.16

39.97 0.01 0.01 13.54 0.24 46.86 0.27 0.00 0.01 0.01 0.17

40.47 0.02 0.02 10.15 0.17 49.26 0.28 0.02 0.01 0.03 0.20

40.61 0.01 0.01 10.15 0.18 49.48 0.28 0.02 0.01 0.03 0.21

40.63 0.01 0.01 10.25 0.19 49.08 0.29 0.01 0.01 0.03 0.20

0.88 5.98

0.89 5.51 0.04

0.89 5.74

0.89 5.76

0.89 6.19

0.89 5.86

0.87 5.99

0.85 5.81

0.89 5.76

0.89 5.85

0.89 5.57

Fo (mol) δ18O (‰) sem (wt%)

Pollena

SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O Cr2O3 NiO

40.04 0.01 0.04 12.76 0.23 45.44 0.33 0.02 0.00 0.01 0.11

40.16 0.01 0.03 12.77 0.25 45.45 0.31 0.01 0.00 0.01 0.14

40.44 0.00 0.04 11.10 0.16 47.46 0.32 0.00 0.00 0.00 0.18

40.91 0.01 0.04 12.37 0.15 47.02 0.34 0.02 0.00 0.00 0.18

40.35 0.00 0.03 10.88 0.23 47.35 0.26 0.01 0.00 0.03 0.11

40.28 0.00 0.03 12.78 0.27 45.45 0.28 0.00 0.00 0.01 0.17

40.22 0.01 0.03 12.31 0.18 47.03 0.33 0.01 0.00 0.00 0.19

39.52 0.02 0.01 12.25 0.22 47.04 0.32 0.00 0.00 0.00 0.19

40.99 0.00 0.00 9.61 0.00 48.62 0.36 0.00 0.00 0.00 0.35

40.85 0.00 0.00 9.91 0.00 48.43 0.44 0.00 0.00 0.00 0.34

41.04 0.00 0.00 9.09 0.00 49.11 0.39 0.00 0.00 0.00 0.26

40.68 0.00 0.00 9.01 0.00 49.20 0.40 0.00 0.00 0.00 0.53

0.86 6.67 0.09

0.86 6.39

0.88 6.18

0.87 6.38

0.88 5.92

0.86 7.03

0.87 6.54

0.87 6.08

0.90 6.04 0.14

0.89 6.52

0.90 6.16

0.90 6.23

Fo (mol) δ18O (‰) sem

AS2f

5. Discussion Distinct initial O-isotope composition of Pompeii and/or Avellino and Pollena eruptions (Fig. 2), and the different slopes of δ 18Ool–Fool covariation trends (Fig. 4) indicate that a high δ18O material (possibly carbonate) interacted with different modalities or at a different degree with the mafic melts. According to the data on primary melts from Southern Italy Quaternary volcanism we can rule out that the δ18O values measured on SV mafic crystals are representative of uncontaminated, primary, mantle-derived compositions. The Fo contents of olivine phenocrysts can be used to trace the chemical evolution of the mafic melts that fed the reservoirs involved in the studied eruptions. The δ 18O values of these early crystallized phases could derive from the following type-mechanisms: – crystallization from a magma (slightly) contaminated by carbonate digestion; – crystallization from an uncontaminated magma followed by diffusive, high temperature solid-state isotopic re-equilibration

of the melt-crystals assemblage during successive magma– carbonate interaction. The solid-state O-isotope diffusion coefficients for olivine and clinopyroxene are in the range of 10− 19 to b10 − 21 (m2/s) at magmatic conditions (Connolly and Muehlenbachs, 1988; Farver, 2010; Ingrin et al., 2001; Ryerson and McKeegan, 1994) thereby the time needed to equilibrate millimeter-size crystals is in the order of 106 yrs, 3 orders of magnitude larger than the assumed residence time of early formed crystals (the average time-life for a magma chamber at SV is not longer than a few thousand years; Morgan et al., 2006; Scaillet et al., 2008). Accordingly, we suggest that minerals crystallized within a primary magma that had been 18O-enriched before that significant differentiation occurred. Two main processes of magma–carbonate interaction could be able to produce a substantial increase in magma δ18O value: – bulk carbonate assimilation in the deep crust; – diffusive, fluid-melt equilibration between a primary magma and a high-δ 18O CO2 flux produced by decarbonation of the crustal basement.

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Table 2 Major (wt.%) and trace (ppm) elements composition of pyroxenes from the studied eruptions. Mean = averaged composition from 2–3 point analyses on the same crystal; Stdev = standard deviation; bdl = below detection limit. wt%

SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O Cr2O3 NiO Wo En Fs δ18O (‰)

Avellino

Pollena

Mean

St dev

52.86 0.40 2.52 4.45 0.10 16.27 23.46 0.12 0.00 0.13 0.03 47.25 45.59 7.16 6.63

0.55 0.09 0.68 0.39 0.02 0.57 0.19 0.02 0.00 0.04 0.02 0.65 1.28 0.65

Trace elements, ppm Isotope Li 7 0.62 Be 9 0.38 B 11 0.77 Sc 45 107.6 Ti 49 2680 V 51 194.8 Cr 53 626.2 Co 59 32.5 Ni 60 126.0 Zn 66 14.07 Rb 85 0.03 Sr 88 79.05 Y 89 9.24 Zr 90 15.77 Nb 93 0.073 Cs 133 0.009 Ba 137 0.132 La 139 2.48 Ce 140 9.63 Pr 141 1.89 Nd 146 11.66 Sm 149 4.07 Eu 151 0.93 Gd 157 3.43 Tb 159 0.42 Dy 163 2.27 Ho 165 0.34 Er 167 0.846 Tm 169 0.089 Yb 173 0.64 Lu 175 0.10 Hf 177 0.75 Ta 181 0.013 Pb 208 0.157 Th 232 0.018 U 238 0.018

Mean 53.51 0.30 1.69 3.61 0.09 16.97 23.50 0.11 0.00 0.14 0.02 46.99 47.22 5.79 6.77

0.60 0.83 1.04 88.3 1626 125.4 1053.6 27.3 134.7 11.02 0.03 65.63 5.67 6.97 0.018 bdl 0.100 1.31 5.34 1.03 6.58 2.16 0.56 2.03 0.23 1.33 0.26 0.347 0.074 0.35 0.06 0.39 bdl 0.139 0.008 0.024

St dev

Mean

St dev

Mean

St dev

0.32 0.02 0.09 0.11 0.01 0.05 0.09 0.01 0.01 0.05 0.02 0.08 0.18 0.16

52.95 0.37 2.18 3.96 0.08 16.53 23.46 0.12 0.00 0.18 0.04 47.27 46.36 6.36 6.31

0.48 0.06 0.36 0.35 0.02 0.34 0.13 0.02 0.00 0.12 0.01 0.16 0.74 0.61

53.38 0.33 1.80 3.85 0.09 17.01 23.19 0.12 0.00 0.12 0.01 46.45 47.39 6.16 6.28

0.37 0.04 0.28 0.05 0.01 0.39 0.14 0.02 0.00 0.02 0.02 0.55 0.65 0.12

0.37 bdl 1.43 90.4 2189 163.7 757.6 31.5 135.1 14.76 bdl 84.40 7.82 14.05 0.041 0.018 0.163 2.12 8.97 1.67 10.18 3.51 0.83 2.92 0.38 2.21 0.34 0.796 0.068 0.49 0.06 0.73 0.008 0.268 0.043 0.014

0.62 0.95 1.46 95.5 1918 141.2 913.3 29.4 129.6 13.88 0.07 68.36 7.14 8.92 0.027 bdl 0.132 1.79 6.78 1.39 8.84 2.80 0.82 2.60 0.30 1.89 0.30 0.698 0.069 0.58 0.05 0.40 bdl 0.206 0.030 0.078

Mean 53.75 0.34 2.12 3.71 0.09 17.00 23.13 0.14 0.01 0.37 0.03 46.50 47.54 5.96 6.25

bdl bdl bdl 87.5 1948 112.1 2691.5 22.6 142.6 11.85 0.07 78.48 7.77 9.07 0.078 bdl 0.126 2.00 7.14 1.43 8.68 2.46 0.69 2.16 0.37 1.35 0.23 0.675 0.077 0.40 0.08 0.55 bdl 0.059 0.056 0.009

Using simple mass balance calculations, and assuming a δ 18O value of 5.51‰ as least contaminated (primary?) olivine at Mt. Vesuvius and a δ 18O value of 25‰ for average local meta-limestones and dolostone (Gilg et al., 2001), the δ 18O values measured in olivine and clinopyroxene (and hence in the tephritic and K-basaltic melts from which they crystallized) would account for a variable carbonate assimilation between 6 and 8%. Contamination of a mafic magma by variable amounts (up to 20 wt.%) of sedimentary carbonate rocks has been suggested to explain the different degree of silica undersaturation, alkali enrichment, and FeO/MgO ratios in the differentiation from shoshonitic basalts to tephrites, and to produce extreme foiditic (alkali-rich and silica-poor) compositions during shallow level magma crystallization (Freda et al., 2008; Iacono Marziano et al., 2007; Mollo et al., 2010). In these experimental runs, the amounts of crystallized clinopyroxene and phlogopite increase proportionally with an increasing fraction of carbonate added to the starting melt, and carbonate digestion proceeds in concert with clinopyroxene crystal-

St dev 0.42 0.05 0.33 0.23 0.01 0.12 0.35 0.03 0.01 0.05 0.02 0.70 0.34 0.36

Mean 55.79 0.23 1.30 2.81 0.08 16.78 23.22 0.16 0.00 0.52 0.03 47.57 47.82 4.62 6.26

0.84 0.86 1.44 60.9 1326 62.3 3529.7 23.9 205.4 11.06 bdl 86.36 4.23 4.63 0.030 bdl bdl 2.15 7.18 1.33 6.68 2.24 0.56 1.59 0.19 1.18 0.19 0.455 0.048 0.14 0.03 0.21 bdl 0.141 0.034 0.011

St dev

Mean

St dev

0.29 0.00 0.04 0.05 0.01 0.06 0.12 0.02 0.00 0.14 0.02 0.13 0.16 0.08

47.61 0.41 1.78 3.64 0.10 13.65 21.50 0.10 0.00 0.10 0.02 49.34 43.84 6.82 6.49

16.12 0.07 0.56 0.24 0.01 1.63 3.59 0.02 0.00 0.12 0.01 1.85 0.88 1.09

0.68 0.21 bdl 92.7 2198 116.8 864.5 26.0 144.8 12.92 0.12 94.45 6.56 14.36 0.026 bdl 0.621 2.64 10.69 1.88 10.53 2.79 0.69 2.84 0.30 1.58 0.30 0.634 0.061 0.49 0.06 0.88 0.004 0.253 0.051 0.019

Mean 54.38 0.36 1.78 3.34 0.08 16.47 23.23 0.15 0.01 0.30 0.02 47.58 46.95 5.47 6.61

0.56 0.48 1.01 90.6 2068 102.9 1912.9 26.3 164.6 13.09 bdl 95.30 6.89 11.40 0.037 bdl 0.164 3.71 12.45 2.16 11.87 3.28 0.85 2.55 0.35 1.85 0.29 0.528 0.095 0.58 0.05 0.71 0.011 0.126 0.047 0.036

St dev 0.45 0.03 0.15 0.15 0.01 0.06 0.24 0.01 0.00 0.04 0.01 0.26 0.13 0.25

Mean 52.64 0.38 1.74 3.52 0.08 17.26 23.45 0.15 0.00 0.15 0.03 46.64 47.76 5.59 6.53

0.74 0.27 1.64 85.9 2041 109.2 1018.1 26.0 161.1 12.01 0.07 83.83 5.56 8.88 0.035 bdl 0.109 1.84 7.34 1.54 9.00 2.80 0.67 2.22 0.30 1.32 0.24 0.542 0.070 0.25 0.05 0.49 0.009 0.231 0.055 bdl

St dev 0.17 0.01 0.19 0.23 0.01 0.24 0.17 0.01 0.00 0.07 0.01 0.11 0.40 0.37

Mean 55.04 0.26 1.53 2.99 0.08 16.68 22.85 0.16 0.01 0.58 0.03 47.15 47.89 4.95 6.80

St dev 0.44 0.09 0.26 0.46 0.02 0.39 0.20 0.02 0.01 0.32 0.02 0.55 0.89 0.77

0.80 0.28 1.48 72.1 1788 105.5 2393.3 24.8 184.2 12.07 bdl 81.38 6.03 9.00 0.040 0.004 0.276 1.85 7.42 1.48 8.67 2.89 0.75 2.34 0.30 1.71 0.24 0.672 0.072 0.52 0.07 0.46 0.006 0.365 0.037 0.006

lization, according to the available MgO in the system. Experiments also show that hyaline glass with rare olivine crystals can be produced only in CaCO3-free runs, and moderate (5 wt.%) CaCO3 addition results in highly crystalline, olivine-free products (Mollo et al., 2010). It could be argued that olivine may still be a stable phase during early stages of Mg-rich carbonates (dolomitic limestones to dolomites) assimilation. These latter crop out in the Vesuvius area (e.g. Iacono-Marziano et al., 2009), and could be a viable contaminant for Vesuvian magmas. However, dolomite assimilation acts to increase the MgO activity in the melt thereby producing high-Fo (N0.90 mol.), low-Ni, and high- 18O olivines, associated with clinopyroxenes, which evolve toward Ca-Tschermak and esseneite components (Gaeta et al., 2009; Peccerillo et al., 2010). These features are not detected in SV mafic products, discarding the hypothesis of an important bulk assimilation of Mg-rich carbonates at depth. Experiments of carbonate contamination of Vesuvius melts are even more stringent, as they claim that at least 10–14 wt.% of

L. Dallai et al. / Earth and Planetary Science Letters 310 (2011) 84–95

89

Table 2 Major (wt.%) and trace (ppm) elements composition of pyroxenes from the studied eruptions. Mean = averaged composition from 2–3 point analyses on the same crystal; Stdev = standard deviation; bdl = below detection limit. wt%

SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O Cr2O3 NiO Wo En Fs δ18O (‰)

Pompei

Pollena

Mean

St dev

Mean

St dev

Mean

St dev

Mean

St dev

St dev

Mean

St dev

52.79 0.39 1.75 3.62 0.07 17.46 23.86 0.13 0.00 0.01 0.01 46.76 47.60 5.65 6.42

0.52 0.04 0.24 0.21 0.02 0.31 0.06 0.02 0.00 0.03 0.01 0.26 0.61 0.36

52.27 0.43 1.95 3.78 0.09 17.10 23.87 0.17 0.01 0.17 0.02 47.08 46.95 5.97 6.04

0.18 0.03 0.21 0.24 0.01 0.28 0.17 0.01 0.00 0.09 0.00 0.36 0.67 0.39

51.86 0.45 1.94 3.84 0.09 17.17 23.73 0.15 0.01 0.09 0.03 46.82 47.13 6.06 6.15

0.41 0.04 0.16 0.30 0.01 0.29 0.17 0.02 0.00 0.07 0.01 0.38 0.68 0.49

53.22 0.43 2.74 4.22 0.09 15.92 23.58 0.17 0.00 0.01

0.21 0.02 0.19 0.07 0.02 0.08 0.04 0.02 0.00 0.01

49.81 0.59 2.93 4.53 0.10 15.05 21.89 0.21 0.00 0.04

3.18 0.51 2.84 2.66 0.07 2.88 0.34 0.14 0.00 0.03

52.67 0.51 2.44 3.89 0.08 16.49 23.85 0.13 0.00 0.14

0.84 0.12 0.62 0.39 0.02 0.42 0.23 0.01 0.01 0.04

52.97 0.40 2.12 3.69 0.09 16.78 23.69 0.13 0.01 0.39

0.16 0.02 0.09 0.08 0.02 0.11 0.10 0.02 0.01 0.12

52.21 0.46 2.26 3.64 0.10 17.16 23.21 0.17 0.00 0.26

0.17 0.02 0.08 0.09 0.03 0.10 0.11 0.01 0.01 0.04

53.50 0.32 1.75 3.45 0.09 16.99 23.17 0.15 0.00 0.29

0.23 0.03 0.37 0.53 0.01 0.48 0.18 0.08 0.00 0.13

53.61 0.26 1.32 2.93 0.07 17.49 23.29 0.16 0.00 0.39

0.21 0.05 0.15 0.20 0.01 0.09 0.14 0.01 0.00 0.10

48.02 45.13 6.86 6.65

0.07 0.19 0.13

47.18 44.86 7.97 6.90

1.66 6.86 5.21

47.80 45.99 6.21 7.47

0.11 0.73 0.65

47.39 46.71 5.91 7.41

0.11 0.13 0.13

46.42 47.75 5.83 6.55

0.15 0.20 0.18

46.74 47.68 5.57 6.57

0.43 1.09 0.87

46.61 48.71 4.68 6.65

0.09 0.35 0.29

Trace elements, ppm Isotope Li 7 bdl Be 9 bdl B 11 bdl Sc 45 96.6 Ti 49 2388 V 51 121.0 Cr 53 98.5 Co 59 23.0 Ni 60 83.4 Zn 66 10.22 Rb 85 bdl Sr 88 84.24 Y 89 6.47 Zr 90 9.20 Nb 93 0.039 Cs 133 bdl Ba 137 bdl La 139 1.77 Ce 140 5.90 Pr 141 1.38 Nd 146 7.72 Sm 149 2.24 Eu 151 0.67 Gd 157 2.36 Tb 159 0.24 Dy 163 1.42 Ho 165 0.22 Er 167 0.571 Tm 169 0.092 Yb 173 0.31 Lu 175 0.05 Hf 177 0.45 Ta 181 bdl Pb 208 0.174 Th 232 0.047 U 238 0.020

0.95 0.69 bdl 93.6 2687 95.3 989.3 22.7 117.7 13.31 0.43 99.87 8.58 14.78 0.091 0.026 0.750 3.65 11.38 2.04 11.75 3.26 0.87 3.06 0.36 1.99 0.35 0.885 0.091 0.39 0.03 0.61 0.018 0.201 0.057 0.047

0.51 1.53 bdl 94.2 2443 128.1 642.2 28.3 130.6 13.67 bdl 88.01 6.61 12.64 0.020 0.011 0.052 2.42 9.69 1.74 9.95 2.78 0.83 2.79 0.31 1.59 0.27 0.605 0.085 0.58 0.07 0.61 0.019 0.253 0.034 0.019

0.92 bdl 1.29 89.5 2531 157.0 112.1 25.7 86.6 16.86 0.157 95.77 8.52 14.95 0.108 0.069 0.580 2.84 11.55 1.84 11.25 3.13 0.77 2.66 0.33 1.72 0.32 0.795 0.101 0.39 0.06 0.54 0.049 0.332 0.126 0.013

Mean

1.96 0.73 1.75 80.8 2198 90.2 249.5 21.5 97.1 11.13 0.125 100.13 7.22 10.95 0.020 bdl 0.625 2.36 8.24 1.70 9.45 2.42 0.82 2.54 0.37 1.81 0.19 0.530 0.051 0.40 0.05 0.62 0.019 0.252 0.017 bdl

0.95 1.35 2.47 100.3 3451 133.5 763.4 24.7 121.3 12.68 bdl 87.74 8.87 21.49 0.035 bdl 0.350 2.48 9.26 2.05 11.73 4.01 1.11 3.55 0.37 2.44 0.31 0.850 0.089 0.55 0.07 0.96 bdl 0.318 0.092 0.016

Mean

1.03 0.52 1.13 87.3 2059 107.5 2140.3 21.8 137.7 13.38 bdl 122.21 9.98 15.49 0.035 bdl 0.209 4.31 14.34 3.20 15.97 4.79 1.28 4.16 0.51 2.42 0.32 0.765 0.085 0.26 0.03 0.79 0.010 0.269 0.078 0.011

St dev

Mean

2.29 0.67 1.62 96.6 2486 104.9 1367.7 23.3 131.3 13.28 0.570 111.53 8.78 16.36 0.070 bdl 0.299 3.77 12.39 2.36 12.55 3.65 0.98 3.03 0.40 2.15 0.28 0.708 0.088 0.38 0.07 0.71 0.024 0.261 0.070 0.039

St dev

Mean

3.27 0.67 1.62 83.2 2181 99.6 1264.9 22.8 123.3 15.10 0.570 109.58 7.73 11.54 0.117 bdl 0.395 3.04 10.09 1.89 10.98 2.97 0.91 2.82 0.34 1.79 0.23 0.652 0.092 0.35 0.06 0.47 0.018 0.349 0.086 0.013

St dev

Mean

St dev

3.07 0.79 1.35 61.4 1595 61.4 2736.9 22.2 158.3 16.71 bdl 90.68 4.94 4.58 bdl 0.025 0.225 2.02 5.88 1.17 5.94 1.75 0.46 1.29 0.22 0.97 0.15 0.475 0.048 0.27 0.02 0.23 bdl 0.141 0.040 bdl

(continued on next page)

carbonate assimilation is needed to pass from K-basaltic to tephritic compositions (Iacono-Marziano et al., 2009). In addition, simple mass balance calculations based on O-isotope data constrain the maximum amount of carbonate assimilation able to explain the observed range of δ 18O to about 7% by weight, lower than that suggested by the results of experimental petrology. Using the software Pele (a PC-hosted program to model the crystallization of silicate liquids based on the MELTS algorithm, able to handle variable processes of carbonate assimilation; Boudreau, 1999) the effects of bulk carbonate assimilation on the chemical and isotopic composition of the magma can be modeled. Results of calculations can be used to quantitatively constrain the amount of carbonate assimilation (Appendix 1 and Table 3). In particular, the observed equilibrium mineral paragenesis of olivine and diopside is not consistent with substantial limestone/dolomite assimilation, which predicts early olivine resorption (olivine is present as a crystallizing phase only for assimilation of less than 5% of carbonate), similar to

what is shown by the experiments (Fig. 5). Massive (higher than 10% by weight) assimilation of carbonate rock by a K-trachybasalt would also result in an important increase of CaO accompanied by a decrease of the SiO2 and MgO content of the contaminated magma, up to concentrations never recorded in natural mafic samples (respectively higher than 15.5% and lower than 45% and 4.7%; see Appendix 1). Another problematic aspect of magma–carbonate assimilation is related to the thermal budget of the process. Thermodynamical constraints on the process of magma–carbonate assimilation calculated using the EC-RAFC worksheet (Bohrson and Spera, 2003; Spera and Bohrson, 2001; and references therein) predict that, starting from a trachybasaltic melt at 1200 °C (a good highly conservative approximation for the liquid temperature) 10 wt.% carbonate assimilation would decrease the initial magma temperature by at least 100 °C (Fig. 6), which contrasts with the temperature of crystallization measured for both olivine and diopside-hosted melt inclusions (Cioni et al., 1998). Parameters used in the modeling are listed in Table 3, while magma

90

L. Dallai et al. / Earth and Planetary Science Letters 310 (2011) 84–95

Table 2 Major (wt.%) and trace (ppm) elements composition of pyroxenes from the studied eruptions. Mean = averaged composition from 2–3 point analyses on the same crystal; Stdev = deviation; bdl = below detection limit. Table 2 standard (continued) Pollena Mean

AS2f St dev

Mean

St dev

Mean

St dev

53.74 0.39 2.08 3.42 0.08 16.57 23.04 0.14 0.00 0.31

0.37 0.03 0.21 0.17 0.01 0.09 0.19 0.01 0.00 0.05

53.17 0.45 2.01 3.32 0.08 16.51 23.88 0.16 0.00 0.18

0.84 0.07 0.44 0.20 0.01 0.35 0.12 0.02 0.00 0.08

55.58 0.44 2.40 3.75 0.09 14.19 23.06 0.13 0.00 0.10

0.40 0.03 0.38 0.46 0.00 0.40 0.10 0.01 0.00 0.01

47.19 47.22 5.59 6.79

0.27 0.24 0.28

48.24 46.40 5.36 6.69

0.30 0.63 0.37

50.35 43.10 6.54

0.30 1.06 0.82

bdl 0.73 1.95 108.0 2738 102.9 1942.9 22.5 138.8 14.30 0.056 116.49 9.69 19.21 0.112 bdl 0.410 4.19 13.38 2.67 15.64 4.31 1.03 3.25 0.44 2.21 0.33 0.825 0.056 0.61 0.08 0.90 0.021 0.177 0.087 0.010

0.74 1.21 1.09 93.2 2808 63.5 1135.0 21.2 139.5 12.12 bdl 205.96 10.29 33.32 0.064 0.026 0.216 6.38 20.04 3.77 21.41 5.04 1.41 4.54 0.54 2.54 0.38 0.940 0.119 0.51 0.08 1.26 0.032 0.163 0.089 bdl

1.04 bdl 85.6 2539 98.7 924.5 23.3 145.7 12.09 0.165 119.43 8.25 15.25 0.128 bdl 0.621 3.48 11.94 2.36 14.57 3.92 1.01 3.92 0.45 1.57 0.29 0.870 0.076 0.66 0.08 0.83 0.010 0.520 0.049 0.013

Mean 52.40 0.30 1.33 2.62 0.05 17.74 23.52 0.14 0.01 0.81 0.02 46.77 49.08 4.15 6.51

1.11 1.26 bdl 69.1 1725 57.1 4589.4 18.7 151.2 8.77 0.256 95.55 4.26 6.90 0.029 bdl 0.207 1.51 5.19 1.21 6.52 2.06 0.59 1.58 0.22 1.21 0.17 0.328 0.039 0.22 0.05 0.37 0.012 0.136 0.033 0.015

St dev 0.49 0.02 0.20 0.12 0.01 0.26 0.27 0.02 0.00 0.20 0.01 0.46 0.53 0.22

Mean 53.78 0.31 0.98 3.48 0.10 17.46 23.38 0.12 0.00 0.24 0.02 46.32 48.14 5.54 6.66

St dev 0.31 0.04 0.22 1.22 0.04 0.52 0.39 0.04 0.00 0.28 0.01 0.70 1.28 1.95

1.42 1.19 1.56 68.9 2118 110.5 708.9 28.6 111.1 18.00 0.328 105.95 8.10 8.26 0.131 0.062 3.060 2.52 7.88 1.79 10.21 3.16 0.82 2.43 0.37 1.67 0.23 0.497 0.078 0.43 0.05 0.27 0.013 0.369 0.035 0.027

and carbonate thermodynamical properties used in the calculations are derived from Bohrson and Spera (2003), Haynes (2010), L'vov (2002), and Wyllie and Boettcher (1969). However, it should be noted that that carbonate assimilation in silicate melt may occur via rapid decomposition and degassing of CO2 rather than full-scale melting (Deegan et al., 2010), thereby implying that the amount of energy required may be different (lower) than predicted by EC-RAFC models, and that the amount of assimilation may be underestimated by model calculations. On the other hand, the consistent δ 18O values of olivine and clinopyroxene measured at SV suggest a homogeneous process of magma contamination. This would be hardly achieved by small degrees of carbonate dissolution, likely resulting into local hyper-calcic melt pockets. It is likely that intra-melt homogenization occurs as higher proportions of carbonate are digested/dissolved, the latter driving melt composition towards more evolved compositions. Therefore, on the basis of the 1) occurrence of large olivine phenocrysts, implying olivine stability in the magma in spite of phase

Mean 53.55 0.35 1.54 2.80 0.06 17.50 23.63 0.14 0.01 0.65 0.03 47.09 48.47 4.44 6.68

bdl 1.21 bdl 84.7 2233 70.0 3057.4 18.6 151.3 9.09 bdl 90.44 5.76 13.12 0.020 bdl 0.244 2.17 6.52 1.45 8.16 2.64 0.68 2.11 0.29 1.17 0.21 0.590 0.025 0.43 0.05 0.71 0.016 0.129 0.044 bdl

St dev 0.59 0.02 0.22 0.06 0.02 0.85 0.26 0.02 0.01 0.22 0.01 1.11 1.18 0.12

Mean 53.42 0.47 1.98 3.38 0.07 16.83 23.28 0.13 0.00 0.44 0.03 47.13 47.41 5.46 7.18

0.95 bdl bdl 103.3 2769 93.6 2268.1 21.9 143.6 9.96 0.045 87.22 7.79 18.48 0.022 bdl 0.288 2.58 8.29 1.92 10.72 2.95 0.87 2.64 0.34 1.76 0.28 0.740 0.084 0.34 0.07 0.91 bdl 0.141 0.034 0.010

St dev 0.93 0.02 0.06 0.15 0.01 0.27 0.17 0.01 0.00 0.03 0.03 0.57 0.46 0.21

Mean 54.16 0.34 1.35 2.85 0.07 17.12 23.68 0.11 0.00 0.43 0.04 47.58 47.86 4.57 6.48

1.14 1.47 2.30 81.8 2030 68.0 2078.4 18.2 124.5 8.25 bdl 85.16 5.78 11.39 0.024 0.037 0.214 2.12 6.64 1.26 8.69 2.18 0.57 1.96 0.26 1.23 0.19 0.440 0.042 0.26 0.03 0.47 0.007 0.148 0.064 0.110

St dev 0.37 0.01 0.04 0.08 0.02 0.09 0.08 0.02 0.00 0.03 0.02 0.14 0.26 0.13

Mean 53.94 0.31 1.01 2.98 0.08 16.60 22.96 0.10 0.02 0.29 0.02 47.43 47.64 4.93 7.59

0.41 0.63 1.14 88.1 1868 60.9 1638.0 20.2 111.9 9.88 0.065 89.65 6.60 8.95 0.033 bdl 0.510 1.91 6.17 1.46 8.78 2.93 0.75 2.62 0.31 1.48 0.25 0.466 0.048 0.28 0.04 0.50 bdl 0.318 bdl 0.023

St dev 2.27 0.01 0.04 0.08 0.01 1.02 0.32 0.01 0.01 0.04 0.01 1.41 1.61 0.24

Mean

St dev

51.76 0.22 1.45 2.81 0.07 16.84 22.76 0.14 0.01 0.75 0.02 46.99 48.37 4.64 6.71

2.59 0.03 0.27 0.13 0.01 0.59 0.44 0.01 0.00 0.20 0.02 0.91 0.96 0.30

bdl 0.69 1.88 71.7 1471 71.4 5754.4 19.5 167.1 11.48 0.116 96.25 5.71 6.19 0.011 bdl 0.325 2.55 8.11 1.73 11.01 2.99 0.79 2.79 0.29 1.36 0.22 0.490 0.047 0.35 0.04 0.30 0.017 0.155 0.024 0.016

resorption, which is expected from carbonate assimilation; 2) near primary chemical composition of the clinopyroxene showing no significant increase in Ca-Tschermak and esseneite components; 3) thermodynamic issues related and energy-constrained model calculations, we consider bulk assimilation as an unlikely process to produce the δ 18O values measured in these crystals, and we favor a process of CO2 fluxing through the melt at depth. 6. The effects of CO2 flux over the δ 18O of primary magmas As an inevitable consequence of interaction between magma and sedimentary carbonate, large amounts of 18O-rich (sedimentaryderived) CO2 are released from the carbonates. Because CO2 is an oxygen-rich carrier and fluid-melt oxygen diffusion is enhanced at magmatic temperature, this flux may diffuse through the magma and eventually re-equilibrate its isotopic composition without inducing other significant compositional changes. Experimental data for O-

L. Dallai et al. / Earth and Planetary Science Letters 310 (2011) 84–95

a

b

Pompeii P. AS2f Pollena Avellino P.

20 OLIVINE

91

10 CLINOPYROXENE

8

Number

15 6 10 4 5 2

0 5.0

5.5

6.0

6.5

7.0

7.5

8.0

18O

0 5.7

6.0

6.3

6.6

6.9

7.2

7.5

7.8

8.1

18O

Fig. 2. Histograms showing the variation of the δ18O values in olivine and clinopyroxene from the investigated eruptions.

isotope equilibrium between CO2 and melilite, basalt and silica glasses predict δ 18O values of CO2 at magmatic conditions in the range of 2– 2.5‰ higher than coexisting glass (Appora et al., 2003; Matthews et al., 1998). Due to the high self-diffusion coefficients of oxygen in basaltic melts (in the range of 10 − 7 to 10 − 8 cm 2 s − 1; Muehlenbachs and Kushiro, 1974; Stolper and Epstein, 1991), isotopic equilibrium is achieved in hours to days in the case of high molar oxygen ratios between gas and melt (10 3 to 10 5). Conversely, the silicate fraction will change negligibly if the ratio between CO2 and melt is low, and the extent of oxygen isotope fractionation is recorded in the δ 18O value of CO2 (Stolper and Epstein, 1991). The SV complex, characterized by the occurrence of a thick carbonate basement, represents an ideal site for thermally-induced CO2 production (Iacono-Marziano et al., 2009; Fig. 7), although a deeper source of non-volcanic CO2 (Frezzotti et al., 2009) cannot be ruled out. It follows that olivine and diopside phenocrysts may have crystallized from magma bodies stalling in deep (more than 8–10 km) reservoirs within a CO2-degassing carbonate basement, as hypothesized on the basis of seismic tomography (Auger et al., 2001; De Natale et al., 2006) and experimental petrology (Scaillet et al., 2008). Oxygen isotopic re-equilibration between the magma and CO2 flux occurred at a pressure not lower than 200 MPa (the pressure

7.6

estimated from volatile measures on melt inclusions hosted in olivine and diopside, after Marianelli et al., 2005), suggesting that CO2 was derived from decarbonation of the deeper portion of the carbonatic basement. In this case, the process of CO2 production could be considered as a general effect related to deep magma generation, transfer and intrusion, possibly unrelated to the specific volume of magma undergoing the flux, and inducing no significant thermal/ compositional changes on it. The amount of available carbonate is high whether compared with the volume of interacting magma, and large amounts of CO2 could be continuously available through time. Conversely, an effect of thermal insulation of carbonates from the magma could be more effective in the shallower reservoirs, where magma can reside for a long time (hundreds to thousands of years), differentiate and directly interact with the hosting carbonates, partially digesting them in some cases. At Mt. Vesuvius, present-day CO2 flux (300 t/day; Iacono-Marziano et al., 2009, and references therein) has δ 18O values varying between +23 and +28‰ (Chiodini et al., 2000), typical of CO2 degassed from a carbonate basement (Fig. 7a). Average magma supply during the last 4 ka of activity at SV has been estimated at 2.7–11.2∗10 9 kg/yr (Scandone et al., 2008). If the present-day CO2 flux (1.1∗108 kg/yr) can

δ 18Ocpx (‰)

7.2

Pollena AS2f

6.8

Pompeii 6.4

Avellino

6.0 18

4‰ 0. = O

δ Ool (‰) 18

5.6 5.2

5.6

6.0

6.4 18

6.8

7.2

Fig. 3. The δ–δ plot correlating the average δ O values of olivine and clinopyroxene from each volcanic eruption. Sample bars refer to the standard deviation of each sample population (st.dev./sq.root n_samples). Symbols are the same of Fig. 1. For Pollena eruptions, the two δ18O values correspond to the averages of two distinct samples.

Fig. 4. The δ18Ool vs. Fool negative co-variation trends in the four investigated eruptions. Olivine crystals from all the selected eruptions vary in a restricted range of chemical composition (from Fo86 to Fo91), with the most primitive crystals (Fo91) measured in samples collected from the AS2f eruption. Fo is the olivine composition calculated as [Mg / (Mg + Fe)]. Symbols are the same of Fig. 1.

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22

Table 3 Parameters used in the EC-RAFC model calculation.

20

Liquidus T magma Initial T magma Specific heat of magma Specific heat of assimilant Specific heat of recharge magma Enthalpy of crystallization of magma Enthalpy of melting of assimilant Enthalpy of crystallization of recharge magma 18 O/16O in magma 18 O/16O in assimilant

a

18

d

c

16

O

1200 °C 1200 °C 1484 J/kg K 1170 J/kg K 1484 J/kg K 396,000 J/kg 360,000 J/kg 396,000 J/kg 5.51 25

Curve a b b c d

b

14

Tla 900 900 900 700 650

Ta0 800 600 600 600 500

Ts 850 850 850 650 620

Teq 900 900 852 700 650

Ma0 1.77 1.18 3.08 2.41 2.28

b1

18

tlm tmo cpm cpa cpr hm ha hr

12 10 8 6 4 700.00

be extrapolated to the past, the influence of such a flux to the supply of mantle-derived magma can be calculated. The result is that the ratio of CO2-derived vs. magma-derived oxygen is in the range 0.02–0.08 wt.% (Fig. 7b). Assuming a δ18O for limestone-derived CO2 of 28‰, the calculated δ18O increase of 0.5 to 1.8‰ matches the measured data for the high-δ18O forsteritic olivine. Considering the magma-CO2 gas system as a whole, in a single frame of time, the increase of δ18O in Mt. Vesuvius mafic melts can be modeled by using a simple mass-balance calculation, in terms of CO2 flux, assuming a fractionation factor similar to that of CO2-melilite melt (Appora et al., 2003). Mass balance calculation requires: 18 i 18 i 18 f xCO2
δ OCO2 þ ð1−xCO2 Þ
δ Oglass ¼ xCO2
δ OCO2
18 f þ ð1−xCO2 Þ δ Oglass

with s = solid phase; ol = olivine; m = melt; cpx = clinopyroxene; g = gas phase; i = initial; f = final. The amount of CO2 (xCO2), required to produce the measured shift οf δ 18O from typical mantle values, is in the range of 3–5%, that is compatible with the estimated ratio of magma supply to CO2 flux. 7. Implications on eruptive activity The variability observed for the δ 18O values of mafic magmas at SV is compatible with a sustained flux of carbonate-derived CO2 through the magma, at PT conditions in equilibrium with the crystallization of

800.00

900.00

1000.00

50 10% CaCO3 assimilation

40

40

Wt %

Cpx Gas

30 20

10

10 0

0

Wt %

10% CaCO3-CaMg(CO3)2 assimilation

10% CaMgCO3 assimilation

Gas Cpx Ol

40

Gas Cpx Ol

30 20

10 0 1200

Wt %

20

Wt %

Gas Cpx Mt-Usp

20

30

1300.00

the olivine and clinopyroxene assemblage. The outcome of this conclusion is that CO2 fluxing through magma may play a significant role in the magmatic processes at SV. In particular, due to its limited solubility in magmatic melts at crustal pressures, externally produced CO2 tends to concentrate into the fluid phase in equilibrium with the magmatic melt. Several effects are possible: 1) forced exsolution of water from previously undersaturated melt. This effect can be very important, as CO2 fluxing throughout the magma induces a decrease in the fugacity of the other volatile species (essentially H2O) in the fluid phase, and a corresponding decrease in their solubility in the magma (Dixon and Stolper, 1995; Papale, 1999). This process may promote the exsolution of H2O from the otherwise undersaturated magma, possibly enhancing the ability of the magma itself to erupt explosively. This appears particularly important in the case of nearly volatile-saturated, small, mafic magma bodies. Looking at the recent activity of SV, this effect could have been very important especially in the last 1500 yrs, characterized

5% CaCO3 assimilation

40

1200.00

Fig. 6. Figure shows the variation of isotopic composition of the magma during the process of assimilation, as a function of magma temperature modeled by EC-RAFC (Bohrson and Spera, 2001). Parameters used in the modeling are listed in Table 3. Ta0 = assimilant initial temperature; Tla = liquidus temperature; Ts = solidus temperatures; Teq = equilibration temperature. The grey box shows the range of δ18O values of diopside and olivine.

50

30

1100.00

Tmagma (°C)

10 1180

1160

1140

Temperature (°C)

1120

1200

1180

1160

1140

1120

0 1100

Temperature (°C)

Fig. 5. Diagrams of the modeled (Pele; Boudreau, 1999) mineral phase abundance during a process of carbonate assimilation and equilibrium crystallization. The different diagrams show that in a process of carbonate assimilation, olivine is stable only after assimilation of N10 wt.% of dolomite. All these experiments produce strongly undersaturated residual melts. Initial magma temperature, 1200 °C; initial carbonate temperature, 600 °C. Initial liquid and assimilant from Cioni et al. (2008), Santacroce et al. (2008). The final temperature (Tf) corresponds to eruptive temperature of typical vesuvian magma calculated by Cioni et al. (1998). Olivine (Fo = 86.64) is a stable phase only for simulations without assimilation.

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Fig. 7. a) Schematic illustration of the mechanism of magma/carbonate interaction at Mt. Vesuvius. b) Graphical representation of the δ18O shift as a function of CO2/magma ratio (wt.%) based on simple mass-balance calculation. Figure is not at scale.

by very frequent eruptions of small magnitude and intensity (Cioni et al., 2008). 2) Overall decrease of the density of a shallow residing magma, due to the introduction of a poorly soluble volatile component like CO2, which could force magma rise by increasing its buoyancy. A similar mechanism has been proposed to explain explosive eruptions of mafic magmas at the Alban Hills Volcano (Freda et al., 2010). 3) If released during local assimilation of the carbonate host rocks in the shallow level magma chamber, CO2 may have different effects according to the size and shape of the reservoir. In fact, the ratio between the volume of the magma and that of the host rocks that exchange heat and mass with the magma (the thermo-metamorphic and metasomatic carapace) is low for small magma chambers. The net result is that smaller is the magma reservoir, larger is the ratio between the mass of carbonate-derived CO2 and magma; thus a large change in the total CO2 fugacity (and consequently H2O solubility) can be imposed on a small magma batch. An important corollary is that the “aptitude” to erupt explosively of the small mafic magma bodies that established at shallow level in the SV area may have been largely increased by local processes of magma–carbonate interaction. We conclude that the O-isotope compositions of the “basaltic” melts at Vesuvius were derived from an early process of CO2 fluxing from the carbonate basement, at the roots of the volcanic structure. This process had the potential to increase the intrinsic explosivity of the mafic magmas feeding the magma chamber. Additional bulk limestone assimilation occurred at shallow depths and mainly involved partially differentiated melts. Supplementary materials related to this article can be found online at doi:10.1016/j.epsl.2011.07.013.

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