Processes and timescales of magma evolution prior to the Campanian Ignimbrite eruption (Campi Flegrei, Italy)

Earth and Planetary Science Letters 306 (2011) 217–228

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Processes and timescales of magma evolution prior to the Campanian Ignimbrite eruption (Campi Flegrei, Italy) Ilenia Arienzo a,⁎, Arnd Heumann b,d, Gerhard Wörner b, Lucia Civetta a,c, Giovanni Orsi a a

Istituto Nazionale di Geofisica e Vulcanologia, sezione di Napoli Osservatorio Vesuviano, Via Diocleziano 328, Napoli, Italy Geowissenschaftliches Zentrun Göttingen, Abteilung Geochemie, Goldschmidtstr. 1, 37077 Göttingen, Germany Dipartimento di Scienze Fisiche, Università di Napoli “Federico II”, Monte S. Angelo, Napoli, Italy d GFZ German Research Centre for Geosciences, Telegrafenberg, 14473 Potsdam, Germany b c

a r t i c l e

i n f o

Article history: Received 17 December 2010 Received in revised form 31 March 2011 Accepted 4 April 2011 Available online 30 April 2011 Edited by: T.M. Harrison Keywords: Campanian Ignimbrite eruption U–Th isotopes residence time mixing process

a b s t r a c t The Campi Flegrei caldera collapsed 39 ka in the Neapolitan area (southern Italy) after the Campanian Ignimbrite eruption. This eruption, recognized as the largest and the most cataclysmic volcanic event in the Mediterranean area over the past 200 ka, extruded not less than 300 km3 of trachytic magma. Controversy exists over the timescales required to assemble such large volume of silicic melt and thus whether large magmatic reservoirs can actually persist below active volcanic systems over prolonged periods of time. Uranium-series analyses have been performed on Campanian Ignimbrite whole-rocks, glass matrixes and separated minerals, and the obtained results have been interpreted in combination with data on Sr, Nd, and Pb isotopes from literature. The compositionally most evolved sample which is most radiogenic with respect to Sr isotopes records a reference age of 71 ka. By contrast, U–Th internal isochrones of the three compositionally least evolved samples give identical initial Th isotope ratios and yield consistent ages predating the eruption by up to 6.4 ka. The highest Pb and Nd isotopic ratios and 230Th/232Th activity ratios together with the oldest reference age of the most evolved samples suggest the existence of a resident magma body possibly related to a magmatic system that is known to have fed earlier magmatic activity in the Campi Flegrei area. Conversely, the younger age of the least evolved and least radiogenic magma dates the crystallization/differentiation event of a chemically and isotopically new magma batch entering the reservoir of the resident magma some few thousand years before the cataclysmic eruption. Therefore, the time preceding this large caldera-forming eruption during which the large volume of Campanian Ignimbrite magma assembled and mixed is 6.4 ± 2.1 ka. © 2011 Elsevier B.V. All rights reserved.

1. Introduction Caldera-forming silicic eruptions are among the most cataclysmic volcanic phenomena documented in the geological record (Smith, 1979). They require generation, accumulation and storage of a large volume of magma in a shallow crustal reservoir for some time prior to eruption. During this time interval magma undergoes crystallization and recharge until eruption begins, likely triggered by arrivals of batches of hotter magma in the reservoir or when the resident magma reaches its level of volatile saturation during crystallization. However, how rapidly a large volume of magma can accumulate, how long it resides prior to eruption, and what the priming time is for an evolved magmatic system to erupt cataclysmically, are still unresolved topics. Magmatic systems may evolve in different ways: some show evidence

⁎ Corresponding author. Tel.: + 39 0816108365; fax: + 39 0816108351. E-mail address: [email protected] (I. Arienzo). 0012-821X/$ – see front matter © 2011 Elsevier B.V. All rights reserved. doi:10.1016/j.epsl.2011.04.002

for prolonged residence time, others for rapid formation or accumulation from pre-existing smaller magma batches. Short-lived U-series isotopes are an important tool in assessing timescales of magmatic processes (Allegre and Condomines, 1976; Condomines et al., 1988; Condomines and Sigmarsoon, 2003; Cooper and Reid, 2008; Gill and Condomines, 1992; Hawkesworth et al., 2004; Reid, 2003). 238U/230Th isotope systematic can potentially provide age constraints on the growth of mineral phases, reflecting crystallization and differentiation processes at timescales of 10– 300 ka. Furthermore, 230Th/232Th activity ratio in volcanic rocks, can be used as an isotopic tracer of the Th/U ratio of the source. Magma residence times have been investigated mostly for caldera-forming eruptions fed by highly differentiated rhyolitic magmas (i.e. Yellowstone and Long Valley calderas — USA; Taupo Volcanic Center — New Zealand; and Olkaria Volcanic Complex — Kenya Rift Valley) (Halliday et al., 1989; Hildreth, 1981; Macdonald et al., 1987, 2008; Sutton et al., 2000). For such large (N500 km3) and oversaturated silicic systems, magma residence times ranging from 10 to more than 100 ka have been inferred (Davies et al., 1994;

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Halliday et al., 1989; Heumann and Davies, 2002; Turner and Costa, 2007; Vazquez and Reid, 2002). For smaller and more mafic alkaline systems, significantly shorter residence times (10–100 years up to a few thousand years) are reported (Bourdon et al., 1994; Hawkesworth et al., 2004; Pyle, 1992; Scheibner et al., 2008; Schmitt et al., 2010; Sigmarsson, 1996). The Campanian Ignimbrite (CI) caldera-forming eruption in the Campi Flegrei volcanic area (southern Italy; Fig. 1a) is intermediate in terms of size, composition and alkalinity between large silicic and small highly alkaline magma systems (Bohrson et al., 2006; Civetta et al., 1997; Fedele et al., 2003; Fedele et al., 2008; Fisher et al., 1993; Fowler et al., 2007; Ort et al., 2003; Pappalardo et al., 2002a; Rosi et al., 1999; Signorelli et al., 1999, 2001). De Vivo et al. (2001) dated this eruption by 40Ar/39Ar at 39.3 ± 0.1 ka. This age was later confirmed by Fedele et al. (2008) who reported plateau ages between 37.1 and

39.5 ka. The error provided on the age determination by De Vivo et al. (2001) represents the internal error. For this reason and acknowledging this limitation on the precision of the eruption age, for the purpose of this paper we will compare our U–Th mineral ages with a reference age for the CI eruption of 39 ka. The oldest subaerial volcanic products in the Campi Flegrei area yield ages of about 60 ka (Pappalardo et al., 1999) and are related to volcanism extending beyond the limits of the present caldera (Orsi et al., 1996). However, volcanic rocks dated at about 150 ka and 80 ka at the neighboring islands of Ischia (Cassignol and Gillot, 1982; Vezzoli, 1988) and Procida (Rosi et al., 1988a,b), respectively, document older long-standing evolved magmatic systems within the Phlegraean Volcanic District (Procida and Ischia islands and Campi Flegrei). However, among all the eruption that occurred in the region, the CI was the most cataclysmic, erupting not less than 300 km3 of

Fig. 1. (a) Structural sketch map of the Campi Flegrei caldera, modify after Orsi et al., 2004; (b) Geographic distribution of the Campanian Ignimbrite deposits, from Fedele et al. (2003); (c) Shadel relief of the Campanian region including the Somma-Vesuvius volcano and the Phlegraean Volcanic District (Campi Flegrei, Procida and Ischia). The quarries were the analyzed samples were collected, are shown (full stars).

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trachy-phonolitic magma. Pyroclastic products of this eruption were deposited at great distance from the vent as shown by the occurrence of ash in late Pleistocene Mediterranean marine sediments and in several archeological sites of South-eastern and North-eastern Europe (Fedele et al., 2003, 2007; Thunnel et al., 1979; Fig. 1b). Atmospheric loading by sulfuric aerosol, ash and CO2 from this eruption may have accelerated cooling of Earth's climate (Fedele et al., 2003, 2007). The CI eruption was accompanied by the collapse of an area of about 230 km2 that was the first expression of the nested Campi Flegrei caldera that evolved further during the past 39 ka (Orsi et al., 1992, 1996; Acocella et al., 2004). After more than 20 ka the second largest collapse within the caldera occurred in relation to the Neapolitan Yellow Tuff eruption (NYT; 14.9 ka ± 0.4 ka; Deino et al., 2004; Orsi et al., 1992, 1995), that extruded N40 km3 of latitic to trachytic magma. Even though such large caldera-forming eruptions are not among the expected events in case of renewal of volcanism in short- to mid-terms (Orsi et al., 2004, 2009), the Campi Flegrei caldera is one of the most hazardous volcanic systems on Earth. This is because of its persistent state of activity and the explosive character of volcanism (Costa et al., 2009; Orsi et al., 1999a,b; 2004). Due to the high volcanic hazard and the large population (1.5 million people) living within the caldera and its surroundings the volcanic risk of this area is extremely high (Orsi et al., 2004). Guided by the previous studies we measured U–Th isotopes by Thermal Ionization Mass Spectrometry (TIMS) on whole rock, glass and phenocrysts from pumice fragments of the CI eruption. The main aims of our study were i) to better constrain the processes leading to the formation of the CI magma chamber, ii) to further characterize distinct magmas involved in the eruption and iii) to place temporal constraints on the pre-eruptive ages and residence times of magmas accumulated prior to this eruption. The new U–Th isotope data are discussed together with our published Sr and Nd isotope data on the same rocks and Sr, Nd and Pb isotopes from literature, in order to constrain the relationships between processes of crystallization/ differentiation and mixing, magma storage and eruption through time. This temporal aspect of magma evolution is extremely important in order to assess the history and behavior of such a hazardous volcanic system. 2. The CI magmatic feeding system The superposition of two compositionally different CI units was only observed in one exposure at Mondragone. However, a core drilled in the northern part of the city of Naples (Pappalardo et al., 2002a) offered the only opportunity to study in a continuous stratigraphic sequence the variable pyroclastic currents deposits emplaced during the CI eruption. The information yielded by the drilled core, together with that from exposed CI deposits, allowed Pappalardo et al. (2002a) to reconstruct the history of the eruption and its relationships with magma withdrawal and caldera collapse. According to this reconstruction the eruption began with phreatomagmatic explosions followed by the formation of a sustained Plinian eruption column fed by the simultaneous emission of two chemically different magmas. Towards the end of this phase, due to upward migration of the fragmentation surface, to reduced magma eruption rate and/or to opening of new fractures, an unstable pulsating column formed and was fed only by the upper, most evolved magma (Pappalardo et al., 2002a). This Plinian phase ended with the waning of the eruption column and the beginning of the caldera collapse. At this stage highly expanded pyroclastic currents were fed by the most evolved magma. During the following major caldera collapse episode, the maximum mass discharge rate was reached and two compositionally different magmas were contemporaneously tapped and mingled during eruption resulting in an intermediate composition hybrid melt. Towards the end of the eruption, only the deepest and

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least differentiated magma was tapped and fed pyroclastic currents that only traveled relatively short distances (Pappalardo et al., 2002a). Geochemical and isotopic investigations on the CI erupted products have been performed by many authors (Arienzo et al., 2009; Bohrson et al., 2006; Civetta et al., 1997; Fedele et al., 2008; Fowler et al., 2007; Pappalardo et al., 2002a; Rosi and Sbrana, 1987; Signorelli et al., 1999, 2001). Civetta et al. (1997) and Pappalardo et al. (2002a) suggested that the CI magma chamber contained two magma layers at variable degree of differentiation and separated by a compositional gap. The upper layer was trachy-phonolitic, while the deeper was trachytic in composition. These magma layers, characterized also by different isotopic compositions, were erupted either separately or simultaneously during the course of the eruption. In this latter case the eruption produced hybrid, compositionally intermediate melt. To explain the different isotopic signature of the CI end-member magmas, Civetta et al. (1997) proposed contamination of the most differentiated part of the magma chamber by radiogenic Sr after feldspar crystallization. Bohrson et al. (2006) suggested assimilation by older partially melted volcanic rocks at depth, whereas Fowler et al. (2007) proposed assimilation of a mixture of hydrothermally altered feldspathoid-syenite and aluminous skarn. The extent of mixing was further discussed by others authors (Signorelli et al., 2001; Webster et al., 2003) mostly on the basis of melt inclusions studies. Signorelli et al. (2001) proposed pre-eruptive magma mixing due to an input of mafic magma into a more evolved system, followed by crystallization of the mixed, hybrid magma toward more evolved compositions. Webster et al. (2003) suggested that the sharp chemical transitions in zoned clinopyroxene phenocrysts and the presence of two populations of melt inclusions (high-MgO and lowMgO) is consistent with mingling and/or mixing between a primitive and a more evolved magma prior to eruption. The authors also suggested that magma mixing/mingling must have occurred shortly before eruption, as the inclusions exhibit no evidence of mixed magma compositions. Arienzo et al. (2009) confirmed the existence of isotopically distinct magma batches and proposed an independent origin for the two CI magmas on the basis of a detailed isotope investigation. Their results showed evidence for partial isotopic disequilibria between the CI magmas and various mineral phases. In the following, we will refer to the most evolved, most radiogenic CI magma as Magma Type 1 (MT 1) and to the least evolved, least radiogenic magma as Magma Type 2 (MT 2). MT 1 is similar in Nd isotopes to the pre-CI magmas (0.51250–0.51252, Table A, Supplementary material; D'Antonio et al., 2007; Pabst et al., 2008; Tonarini et al., 2004). Conversely, MT 2, characterized by 143Nd/144Nd ≈ 0.51248 (Table A, Supplementary material), is distinct from the compositions recorded in pre-CI magmatic activity in the area and thus was never erupted at Campi Flegrei prior to the CI eruption. Arienzo et al. (2009) suggested that these two isotopically distinct magmatic components “met” and mixed before onset of the CI eruption, being not genetically related by processes as crystal fractionation and assimilation, and interpreted the evolved end-member as a resident magma, whereas the more mafic end-member was considered to represent new magma entering the reservoir before the eruption. 3. Sampling strategy For the purpose of this study and according to the chemostratigraphic reconstruction of Civetta et al. (1997) and Pappalardo et al. (2002a), samples representative of each of the different magmas erupted during the CI eruption were collected. MT 1 and the intermediate composition hybrid magma were sampled at Mondragone, along the South-eastern slopes of Monte Massico, 38 km from the vent area (Fig. 1c). The composite pumice sample Mondragone

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152a2 (Mond152a2), representative of MT 1, was collected from the lower unit, whereas the composite pumice sample Mondragone 15U3 (Mond15U3), representative of the compositionally intermediate, hybrid magma, was sampled from the upper unit. MT 2 was sampled at San Marco Evangelista (San Marco E., Fig. 1c) in a quarry along the westernmost slopes of the Appennine Mountains, about 30 km from the vent area. At the same stratigraphic level, a composite pumice sample (constituted by black and light gray pumice fragments) as well as a single, large (≈30 cm in diameter), black pumice fragment were collected. The composite pumice sample was split, according to the color of the pumice components, into a black and a light gray fraction, then separately analyzed. The three samples from San Marco Evangelista were labeled San Marco OF 3D black pumice (SMs), San Marco OF 3D black single pumice (SMsp) and San Marco OF 3D light pumice (SMc). To shed light on the processes that occurred within the Campi Flegrei magmatic feeding system before the formation of the CI magma reservoir, we performed uranium-series analyses on two additional samples that predate the CI eruption (58 ka and 47 ka; Pappalardo et al., 1999). 4. Analytical methods Pumice samples were washed several times in de-ionized water, then dried, crushed to particles b1 cm and powdered in an agate mortar. The outer part of each pumice fragment was removed before preparation, by using a dentist drill. Whole rocks were analyzed for major and trace elements (see Arienzo et al., 2009 for details) at the Georg August Universität, Geowissenschaftliches Zentrun Göttingen (GZG), Göttingen (Germany) by X-ray fluorescence spectrometry (XRF) and Inductively Coupled Plasma Mass Spectrometry (ICPMS). The analyzed samples are trachytic in composition and show slightly different degrees of evolution (SiO2 from 59.5 to 61.5 wt.%; Zr from 279 to 689 ppm). In the MgO (wt.%) vs. Differentiation Index (DI, Thornton and Tuttle, 1960) plot, they fall within the range of previously analyzed CI compositions (Fig. 2) and are representative of the different CI magmatic components. More details on chemical variations in the CI erupted products are discussed in Arienzo et al. (2009), Bohrson et al. (2006), Civetta et al. (1997), Fedele et al. (2008), Fowler et al. (2007), Pappalardo et al. (2002a), and Signorelli et al. (1999). The analyzed pumice fragments are glassy (70–97 vol.%) and highly vesicular. The least evolved fragments contain ≈10–30% by

volume of phenocrysts, while the most evolved generally contain b3%. In both cases, phenocrysts are dominantly alkali feldspars with subordinate plagioclase, Fe and Mg-rich pyroxenes, biotite and magnetite. Apatite is an accessory phase. Feldspars were hand-picked from the cleanest fraction (b0.5 mm). Pyroxene was distinguished under the binocular microscope by color and separated by hand picking into dark Fe-rich and light Mg-rich crystals. Magnetite was separated by means of a hand magnet. Phenocrysts were cleaned under the binocular by discarding grains with glass rinds attached or other minerals included. After inspection for mineral inclusions, about 100 mg of glass and 300 to 500 mg of mineral grains were washed with ultra-clean water in an ultrasonic bath to remove remaining glass. Whenever possible, different sized phenocrysts were separated by hand picking (b0.5 mm, ~0.5 mm and 0.6–1.5 mm). The separated phases were dissolved in pressurized Teflon beakers using a mixture of HF–HNO3–HClO4, then dried down and repeatedly converted into nitrate and hydrochloric form. Subsequently sample solutions were spiked with tracer solutions enriched in 229Th and 236U. Chromatographic separation of U and Th from the solutions followed a standard anionic scheme on miniaturized columns with total procedural blanks well below 15 pg equivalent to less than 0.01% of the total element fractions analyzed. The alkaline earths and REE-bearing eluants of this separation scheme were used for Sr and Nd isotope analyses. Results are in Arienzo et al. (2009). U and Th concentrations were determined by isotope dilution. U and Th isotopic ratios were measured by using the MAT262 RPQ+ thermal ionization mass-spectrometer at the GZG. Depending on ion beam intensities, 232Th/229Th, 234U/238U and 236U/238U isotope ratios were analyzed either statically, using the ion counting system and a Faraday detector, or only with the ion-counter, by switching the magnet settings. U-isotope ratio measurements were processed offline following an exponential law of mass fractionation after correction for the added tracer solution. Repeated analyses of U112a (CRM-960) gave a 234U/238U = 0.0000524 ± 0.0000003 (2σ; n = 10) with 50–100 ng U loaded on the filaments. External precision of Th isotope ratios was monitored with an in-house standard, with a 230 Th/232Th within the range of the analyzed silicic samples at ±0.7% (2σ). The reproducibility and accuracy of U and Th isotope ratio measurements was monitored during the course of this study by isotope analyses of an international accepted standard (Table Mountain Latite, TML; e.g. Goldstein et al., 2003). U–Th data for the analyzed fractions are reported in Table 1, whereas Sr and Nd isotopic compositions (from Arienzo et al., 2009) are listed in Table A (Supplementary material). Isochrones were calculated via linear regression based on McIntyre et al. (1966), Titterington and Halliday (1979), and York (1969). All errors are evaluated at 2σ confidence level. 5. Results and discussion 5.1. U–Th disequilibria

Fig. 2. Chemical variability of the CI deposit. The reference field is built on literature data from Civetta et al. (1997), Fedele et al. (2008), Pappalardo et al. (2002a), and Signorelli et al. (1999). The compositions of the analyzed samples, representative of the CI MT 1 and MT 2, cover a large part of the compositional range displayed by the CI products.

Measured 238U/234U activity ratios (Table 1) on CI whole rocks, glasses and phenocrysts of selected samples range from 0.996 ± 0.011 to 1.017 ± 0.020, suggesting that the analyzed samples have not been affected by secondary alteration. Positive linear relationships exist between degree of differentiation and Th and U concentrations as well as 230Th/232Th activity ratios, as well as Sr and Nd isotope ratios (Fig. 3). U–Th disequilibria for whole rocks and glasses are shown in the equiline diagram of Fig. 4. The CI whole rocks have U/Th activity ratios (hereafter indicated in parentheses) lower than their matrix glasses and all fractions are characterized, compared to secular equilibrium, by an excess (238U) relative to (230Th) ranging from 7% to 10%, the highest values are recorded in glasses from the

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Table 1 U–Th data for the analyzed samples from the Campanian Ignimbrite eruption. Samples

Sample type

[Th] (ppm)

±2σ

[U] (ppm)

±2σ

(238U/232Th)

±2σ

(230Th/232Th)

± 2σ

(230Th/238U)

TML 1 TML 2 SMsp SMsp SMsp SMsp SMspa SMsp SMsp SMsp SMsp SMsp SMsp SMs SMs SMs SMs SMs SMs SMs SMs SMs SMc SMc SMc SMca SMc SMc SMc SMc SMc Mond15U3 Mond15U3 Mond15U3 Mond15U3 Mond152a2 Mond152a2 Mond152a2 Mond152a2a 9601A1 (58 ka) 947 (47 ka)

wr wr wr wr Glass mt mt cpx (0.6–1.5 mm) cpx b 0.5 mm cpx (~0.5 mm) bt (0.6–1.5 mm) bt N 1.5 mm Feld b 0.5 mm wr Glass mt cpx (0.6–1.5 mm) cpx (0.5 mm) cpx b 0.5 mm bt (0.6–1.5 mm) bt b 0.5 mm Feld b 0.5 mm wr Glass mt mt cpx (0.5 mm) cpx (0.6–1.5 mm) bt (0.5 mm) bt b 0.5 mm bt (0.6–1.5 mm) wr Glass cpx (0.5 mm) mt wr Glass cpx (0.6–1.5 mm) cpx (0.6–1.5 mm) wr wr

30.16 29.96 21.44 21.76 34.13 3.102 3.181 2.625 2.368 2.756 1.946 1.633 0.058 22.55 34.19 4.094 2.377 2.087 1.867 1.102 1.249 0.283 40.58 36.86 3.597 3.653 2.572 3.216 1.573 1.897 1.510 48.01 48.53 4.346 5.439 51.45 44.08 4.581 5.019 21.11 48.40

0.29 0.11 0.06 0.11 0.08 0.013 0.012 0.007 0.007 0.018 0.007 0.003 0.001 0.05 0.17 0.008 0.012 0.014 0.006 0.003 0.005 0.003 0.09 0.07 0.007 0.015 0.005 0.014 0.016 0.004 0.005 0.34 0.44 0.035 0.022 0.10 0.13 0.017 0.075 0.05 0.13

10.37 10.43 7.00 7.01 11.46 0.492 0.518 0.516 0.536 0.550 0.354 0.431 0.019 7.38 11.53 1.520 0.586 0.422 0.398 0.245 0.277 0.093 13.32 12.52

0.05 0.10 0.02 0.06 0.07 0.014 0.004 0.006 0.004 0.003 0.003 0.022 0.001 0.09 0.09 0.046 0.004 0.002 0.002 0.001 0.001 0.002 0.39 0.04

1.044 1.056 0.991

0.011 0.011 0.004

1.043 1.037 0.914

0.041 0.010 0.008

0.999 0.982 0.922

1.019

0.006

0.932

0.007

0.915

0.494 0.596 0.687 0.605 0.552 0.801 0.981 0.992 1.024 1.126 0.749 0.614 0.647 0.675 0.672 0.994 0.996 1.030

0.004 0.007 0.006 0.005 0.005 0.040 0.025 0.013 0.010 0.034 0.006 0.005 0.003 0.003 0.004 0.020 0.029 0.004

0.759 0.789 0.808 0.804 0.774 0.792 0.904 0.925 0.932 0.946 0.807 0.778 0.786 0.798 0.800 0.878 0.927 0.933

0.014 0.008 0.010 0.016 0.015 0.007 0.099 0.006 0.011 0.015 0.016 0.012 0.010 0.010 0.016 0.053 0.008 0.007

1.536 1.324 1.176 1.329 1.402 0.989 0.922 0.932 0.910 0.840 1.077 1.267 1.215 1.182 1.190 0.883 0.931 0.906

0.712 0.495 0.777 0.315 0.391 0.275 16.81 16.52 1.132 1.038 17.3 15.42

0.009 0.002 0.004 0.001 0.002 0.001 0.16 0.08 0.008 0.007 0.16 0.08

0.591 0.584 0.734 0.607 0.625 0.552 1.062 1.033 0.791 0.579 1.020 1.062

0.008 0.003 0.005 0.006 0.004 0.003 0.012 0.011 0.008 0.005 0.010 0.006

0.784 0.783 0.836 0.796 0.792 0.768 0.942 0.947 0.855 0.808 0.937 0.960 0.896

0.011 0.008 0.011 0.032 0.007 0.014 0.020 0.019 0.022 0.011 0.007 0.009 0.010

1.327 1.341 1.139 1.311 1.267 1.391 0.887 0.917 1.081 1.396 0.919 0.904

1.277 6.33 16.30

0.006 0.04 0.16

0.772 0.910 1.022

0.012 0.006 0.010

0.853 0.938

0.008 0.009

0.937 0.918

a

(238U/234U)

± 2σ

0.997

0.016

1.007 1.002 1.017

0.013 0.015 0.020

1.010 1.001

0.029 0.003

1.007

0.012

1.003 0.996

0.009 0.011

1.005 0.999

0.009 0.018

1.022

0.016

Different aliquot of minerals separated from the same sample.

Mondragone samples ((238U/230Th) ~ 1.11). Such samples represent the CI MT 1 characterized by 87Sr/86Sr ~ 0.70740, Nd isotopes similar to the pre-CI magmas (~0.51251) and 206Pb/204Pb of 19.19. MT 2, having (238U/230Th) of ~ 1.08, differs from the previous one by lower 87Sr/86Sr (~0.70730), 143Nd/144Nd (~0.51248) and 206Pb/204Pb (19.12). The majority of mineral phases have lower (238U/232Th) and 230 ( Th/232Th) than whole rocks and glasses (Table 1) so that a large spread of U/Th ratios is obtained, providing for a relatively good age resolution. The Mg-rich pyroxenes have been interpreted as xenocrysts within the CI magmas on the basis of Sr and Nd isotope compositions (Arienzo et al., 2009) as well as of clinopyroxene/liquid trace element partitioning considerations (Fedele et al., 2009). Their low Th and U contents have not permitted measurement of the activity ratios. The largest fraction of biotite crystals from the SMsp sample lies – within errors – on the equiline being characterized by (230Th/232Th) of 0.792 ± 0.007 and (238U/232Th) of 0.801 ± 0.040 (Table 1). Mg-rich pyroxene and this biotite fraction possibly represent either xenocrysts incorporated from the wall rocks or crystals recycled from earlier phases of magmatism. Excluding them, the U–Th concentrations measured on minerals and glasses, the latter assumed as those of the melts, can be used to evaluate mineral/melt partition coefficients (Kds) relative to the least differentiated samples (Table 2). Generally KdTh is higher than the KdU for individual minerals, with the exception of the analyzed feldspar fractions which

have KdTh equal to KdU (from 0.002 to 0.008). This characteristic is confirmed by the values reported in literature for plagioclase, for which KdTh and KdU were both in the range 0.01–0.03 for similar rocks (e.g. Villemant, 1988; Wörner et al., 1983). However, the KdTh and KdU reported in this study for feldspar are one order of magnitude lower than those obtained on feldspars from compositionally similar rocks (Geochemical Earth Reference Model, GERM, database). Our lower values confirm that a careful picking may have avoided crystals containing mineral and/or melt inclusions. The biotite fractions have KdTh from 0.032 to 0.057 and KdU from 0.021 to 0.038. These values, and those calculated for the magnetite fractions (KdTh ~ 0.09 and KdU ~ 0.05), are well below, or close to the minimum values reported in literature for phenocrysts from Campi Flegrei trachytes (Table 2). Pyroxene has Th and U partition coefficient from 0.05 to 0.09 and from 0.035 to 0.062, respectively, and KdTh/KdU ~ 1.58. The calculated partition coefficients match the lowest values reported by Villemant (1988), but are higher than those reported by Fedele et al. (2009) for pyroxenes in equilibrium with the trachytic CI melt. We explain such differences by the presence in pyroxenes of Th- and U-rich inclusions, such as apatite and melt inclusions. Therefore, Kd values for pyroxene must be considered to be maximum values. However, since our KdTh/KdU ratios (~ 1.58) are similar to those of Fedele et al. (2009) (Kd Th /Kd U ~ 1.54), we conclude that the pyroxenes from the trachytic MT 2 have grown in equilibrium

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a

b

c

d

Fig. 3. U/Th concentration ratios, (230Th/232Th) and Sr and Nd isotope ratios vs. Differentiation Index (DI). Symbols are the same as in Fig. 2.

with their host-melt. The KdTh and KdU obtained for the other minerals, which are also in the lowest range of values reported in literature, then also possibly represent equilibrium partitioning. This

is an essential conclusion and condition for the interpretation of internal mineral isochrones. 5.2. Magma crystallization and residence times

Fig. 4. (230Th/232Th) vs. (238U/232Th) equiline diagram. The activity ratios measured for the CI whole rocks and glasses are shown.

5.2.1. Residence time of the least radiogenic and least evolved magma If indeed phenocrysts from the San Marco samples crystallized from MT 2, we can use their activity ratios to constrain the crystallization age of this end-member. The isochron regression for each sample is based on activity data of only those minerals that have been shown to be in Sr isotope equilibrium with the host glass (Arienzo et al., 2009). This important independent control on equilibrium, together with the previous considerations based on the calculated U–Th partition coefficients, assures that for the age calculation we only used those minerals that are likely to have grown together from the same melt. The occurrence of mineral/melt inclusions within the analyzed phenocrysts, clearly documented by pyroxene for which detailed studies on trace element partitioning already exist, will not affect the age significance, since mixing between different phases of the same age and from the same magma will only affect the spread of the array in terms of U/Th ratios but not the slope of the isochron. As discussed in Section 3, the SMc, SMsp and SMs samples represent the compositionally least evolved and least radiogenic CI MT 2 at the top of the stratigraphic sequence (Pappalardo et al., 2002a). The internal isochrones calculated for these three samples yield ages of 45.9 ± 2.9 ka, 44.5± 3.3 ka and 47 ± 11 ka (Fig. 5a, b, and c). The uncertainties are 2σ. The Mean Square Weighted Deviates (MSWD), calculated using

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Table 2 Calculated partition coefficients are compared with literature data. Sample

Phase

Calculated KdTh

Calculated KdU

KdTh/KdU

SMsp SMsp SMsp** SMsp SMsp SMsp SMsp SMsp SMsp SMs SMs SMs SMs SMs SMs SMs SMs SMc SMc SMc** SMc SMc SMc SMc SMc

Glass mt mt cpx (0.6–1.5 mm) cpx (b 0.5 mm) cpx (~0.5 mm) bt (0.6–1.5 mm) bt (N 1.5 mm) Feld (b 0.5 mm) Glass mt cpx (0.6–1.5 mm) cpx (~0.5 mm) cpx (b 0.5 mm) bt (0.6–1.5 mm) bt (b 0.5 mm) Feld (b 0.5 mm) Glass mt mt cpx (~0.5 mm) cpx (0.6–1.5 mm) bt (0.5 mm) bt (b 0.5 mm) bt (0.6–1.5 mm)

1.000 0.091 0.093 0.077 0.069 0.081 0.057 0.048 0.002 1.000 0.120 0.070 0.061 0.055 0.032 0.037 0.008 1.000 0.098 0.099 0.070 0.087 0.043 0.051 0.041

1.000 0.043 0.045 0.045 0.047 0.048 0.031 0.038 0.002 1.000 0.132 0.051 0.037 0.035 0.021 0.024 0.008 1.000

1.00 2.12 2.06 1.71 1.48 1.68 1.85 1.27 1.04 1.00 0.91 1.37 1.67 1.58 1.52 1.52 1.03 1.00

0.057 0.040 0.062 0.025 0.031 0.022

1.74 1.76 1.40 1.70 1.65 1.87

KdTh from literature

KdU from literature

0.27 0.27 0.05–0.24 0.05–0.24 0.05–0.24 0.09–0.11 0.09–0.11 0.01–0.03

0.29 0.29 0.04–0.17 0.04–0.17 0.04–0.17 0.08 0.08 0.01–0.03

Lemarchand et al. 1987 Lemarchand et al. 1987 Villemant, 1988 Villemant, 1988 Villemant, 1988 Villemant, 1988 Villemant, 1988 Villemant, 1988

0.27 0.05–0.24 0.05–0.24 0.05–0.24 0.09–0.11 0.09–0.11 0.01–0.03

0.29 0.04–0.17 0.04–0.17 0.04–0.17 0.08 0.08 0.01–0.03

Lemarchand et al. 1987 Villemant, 1988 Villemant, 1988 Villemant, 1988 Villemant, 1988 Villemant, 1988 Villemant, 1988

0.27 0.27 0.05–0.24 0.05–0.24 0.09–0.11 0.09–0.11 0.09–0.11

0.29 0.04–0.17 0.04–0.17 0.08 0.08 0.08

Lemarchand et al. 1987 Lemarchand et al. 1987 Villemant, 1988 Villemant, 1988 Villemant, 1988 Villemant, 1988 Villemant, 1988

KdTh from Fedele et al., 2009

KdU from Fedele et al., 2009

0.014 0.014 0.014

0.009 0.009 0.009

0.014 0.014 0.014

0.009 0.009 0.009

0.014 0.014

0.009 0.009

** Different aliquot of minerals from the same sample.

eight measurements for each sample, are 0.38, 1.6, and 2.4, respectively. The weighted mean age is 45.4 ± 2.1 ka and the (230/Th/232Th) initial ratios are identical within errors (SMs = 0.869 ± 0.050; SMsp = 0.883 ± 0.054; SMc = 0.884 ± 0.049) with a weighted mean value of 0.879 ± 0.029. Chemical and isotopic equilibrium between crystals and liquid is a necessary, but insufficient, condition for demonstrating coeval crystallization from the host liquid, and therefore for interpreting crystal ages in terms of magma residence time. However, the occurrence of ages identical (within error) for all three MT 2 samples suggests that they indeed represent the average crystallization/ differentiation age of this magma. This is confirmed by the fact that the ages calculated for each least evolved CI sample, when based on differently sized phenocrysts (large and small fractions), are identical in all the calculations (Table B, Supplementary material). The “partial”, internal mineral isochrones yield ages of ~44 ka (SMsp), ~ 45 ka (SMc) and ~49 ka (SMs), giving further evidence that the minerals formed in equilibrium with the melt shortly before eruption. As all analyzed minerals may or may not contain inclusions of accessory phases, we cannot rule out that our two point isochrones are partly controlled by these minerals. In such a case, the age given by the slope represents a mixing array defined by variable proportions of accessory phases (Heumann and Davies, 2002). The age would be the crystallization age of such accessory phases (mostly apatite and magnetite), which again would have to have crystallized shortly before the eruption onset. Therefore, we interpret the weighted mean age of the MT 2 magma as a crystallization/differentiation event of relatively short duration that also provide a minimum residence time for the magma. Such a time for MT 2, characterized by 143Nd/144Nd of ~ 0.51248, 87Sr/86Sr of ~ 0.70730 (Arienzo et al., 2009), 206Pb/204Pb ~ 19.12 (D'Antonio et al., 2007) and initial Th of 0.879 ± 0.029, is calculated at 6.4 ± 2.1 ka. 5.2.2. Residence time of the most evolved CI magma and evolution of the pre-CI magmatic feeding system Civetta et al. (1997) and Arienzo et al. (2009), on the basis of the bimodality in the glass shards composition and the occurrence of minerals in Sr isotope disequilibrium with the host glass, considered Mond 15U3 sample representative of a compositionally intermediate

magma which formed by mixing of an almost crystal-free most evolved trachy-phonolitic magma (MT 1), here represented by sample Mond152a2, and a crystal-rich, slightly more mafic trachytic melt (MT 2). For this study, we separated ~ 300 mg of Fe-rich pyroxene which is in Sr isotope equilibrium with Mond152a2 whole rock and its matrix glass to measure its Th activity ratio. A second aliquot was prepared for U–Th analysis. We measured the U fraction, although we miss the (230Th/232Th). However, this replicate differs in terms of Th concentration (5.02 ppm) with respect to the first aliquot (4.58 ppm), and was found to be not in Sr isotope equilibrium with the host glass (see Table 1 and Table A). These differences reflect sample heterogeneities, well documented in both, the CI compositionally intermediate and evolved samples (Arienzo et al., 2009; Civetta et al., 1997). For these reasons we could not calculate an internal mineral isochron for the evolved CI samples. However, by using glasses and whole rocks from both Mondragone samples we calculated a “reference” age of 71 ± 58 ka (MSWD = 1.3). If we consider that the large error is only due to the small data spread and/or to the large errors on the Mond15U3 analyzed fractions, and not to a systematic problem in the approach, then this age information suggests that the CI MT 1 perhaps represents an older resident magma with respect to the least evolved MT 2. In Fig. 6 the age-corrected activity ratios of the pre-CI samples are plotted on the equiline diagram together with those of the CI samples. The trachy-phonolitic sample labeled 9601 A1 (58 ka) has Sr and Nd isotope compositions of 0.70686 and 0.51251 respectively, similar to the typical values characterizing the pre-CI magmas (D'Antonio et al., 2007; Pabst et al., 2008; Tonarini et al., 2004). It is also characterized by 206Pb/204Pb of 19.16, and by the lowest (230Th/232Th) activity ratio with respect to all the analyzed samples ((230Th/232Th) = 0.813). The pre-CI sample 947 (47 ka) has Sr isotope composition of 0.70699, Pb of 19.23 and is similar in terms of Nd isotope ratio, (230Th/232Th), and (238U/232Th) activity ratios compared to the CI evolved samples (Fig. 6). In Fig. 6, isotope compositions and activity ratios of the pre-CI and the CI samples are compared with two samples from Procida Island, labeled FAM and UPFUC (Avanzinelli et al., 2008). The old dated Procida

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ac

a

b Fig. 6. (230Th/232Th) vs. (238U/232Th) equiline diagram. The activity ratios of the CI whole rocks and glasses are compared with those of the pre-CI samples dated 47 ka and 55 ka (Pappalardo et al., 1999), of the Procida samples dated 22 ka and 55 ka (Avanzinelli et al., 2008) and with those relative to the post-CI sample and to the continental crust. All activity ratios are age-corrected.

c

Fig. 5. Internal mineral isochrons relative to the least-evolved CI samples. The uncertainties are 2σ.

sample (FAM, 55 ka) has (230Th/232Th) similar to the oldest Campi Flegrei sample (9601 A1), whereas the young Procida sample, dated at 21 ka (labeled UPFUC), has 87Sr/86Sr, 143Nd/144Nd and (230Th/232Th) similar to the CI MT 2. 6. Discussions 6.1. Magma source characteristics The isotopic and chemical variations through time of the Campi Flegrei magmas (Arienzo et al., 2009; D'Antonio et al. 2007; Di Renzo et al., 2011; Pabst et al., 2008; Pappalardo et al., 1999; Tonarini et al.

2004), in particular those erupted in the 60–10 ka time interval, provide useful information on both processes and magma components within the magmatic system feeding Campi Flegrei volcanism. In the following, we will combine our U–Th isotope data and mineral ages with Sr, Nd, and Pb isotope data from literature in an attempt to discuss the Campi Flegrei magma source characteristics. Sr, Nd, and Pb isotope compositions and age corrected (230Th/ 232 Th) of the magmas erupted in the time interval 60–55 ka at Procida, Ischia, and Campi Flegrei range from 0.7051 to 0.7073, 0.51260 to 0.51251, 18.97 to 19.25, and 0.81 to 0.84, respectively (Avanzinelli et al., 2008; D'Antonio et al., 2007, Pabst et al., 2008; Pappalardo et al., 1999; Tonarini et al., 2004). The 238U/230Th ratios are b1.07. Magmas erupted in the 47–39 ka interval, including the CI MT 1, are characterized by 143Nd/144Nd ~ 0.51251, 206Pb/204Pb ~ 19.20, Sr isotope composition ranging from 0.7070 to 0.7074 (D'Antonio et al., 2007; Di Renzo et al., 2011, Pabst et al., 2008; Pappalardo et al., 1999; Tonarini et al., 2004) and age corrected (230Th/232Th) of ~ 0.89. The isotopic signatures displayed by magmas erupted in these time intervals were attributed to modification of a TransitionalMORB-type asthenospheric mantle source by slab-derived fluids and melts. D'Antonio et al. (2007) and Tonarini et al. (2004) explained the stronger enrichments in fluid mobile elements and B isotopes shown by the CI with respect to the pre-CI magmas, as due to the composition and amount of enriching components from the slab. Since the content of fluid immobile elements (e.g. Nb, Zr, and Th) is relatively high (but still depleted with respect to LILE) the authors also suggested involvement of sediment partial melts. This enrichment process explains the isotopic features of the pre-CI labeled 947 and of the CI most and least evolved samples, with respect to the pre-CI 9601 A1 (58 ka), i.e. the increase of (230Th/232Th) from ~0.81 to ~0.89, 238U/ 230 Th from ~1.07 to ~1.11 and 206Pb/204Pb from 19.16 to 19.23, at fairly constant Nd isotope compositions in the 60–39 ka time interval. However, we have to notice that the CI MT 2 is characterized by 87 Sr/86Sr of 0.7073, lower Nd and Pb isotope ratios of ~0.51248, and 19.12 (Arienzo et al., 2009; D'Antonio et al., 2007; Di Renzo et al., 2011) and a lower (230Th/232Th) ratio of ~ 0.88 compared to both CI MT 1 and the 47 ka old pre-CI magma. Furthermore, CI MT 2 and some of the post-CI erupted magmas, also characterized by low Nd and Pb and variably high Sr isotope ratios, in terms of U-series isotopes, trend towards the least differentiated and most contaminated magma erupted ~ 10 ka ago at Campi Flegrei: Minopoli 2 shoshonite

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(Pappalardo et al., 2002b), having U–Th activity ratios similar to those of the continental crust (Fig. 6). The isotopic features displayed by these post-CI magmas as well as MT 2 are consistent with the involvement of a continental crust component in the evolution history of some of the Campi Flegrei magmas.

225

a

6.2. Genesis and differentiation of the CI magmas Primary magmas at Campi Flegrei, derived from the asthenospheric mantle modified by slab-derived fluids and melts, stagnate in deep reservoir/s with the top at ~ 8 km depth. At this level magmas differentiate and then rise to shallower reservoirs at ~ 5 km depth where they further differentiate and mix/mingle. Such a picture of the structure of the Campi Flegrei system is mostly based on the results of both melt inclusions studies on CI (Marianelli et al., 2006) and younger than 10 ka Campi Flegrei erupted products (Mangiacapra et al., 2008; Arienzo et al., 2010) and seismic tomography investigations (Zollo et al., 2008). According to this model, and on the basis of our isotope studies, we suggest that before the CI eruption poorly evolved magmas were stored in the deep reservoir (~ 8 km), slightly differentiated and rose to shallower depth (~ 5 km) forming discrete magma batches (Pabst et al., 2008), partially preserving their Nd and Pb isotope compositions. These magmas further differentiated and fed some of the pre-CI eruptions. CI MT 1 likely derived from one of these magma batches. Arienzo et al. (2009) suggested that a resident pre-CI magma evolved at shallow depth by combined fractional crystallization/assimilation of sediments enriched in radiogenic Sr, increasing its Sr isotope composition to reach the chemical and isotopic composition of the CI MT 1, without changing 143Nd/144Nd because of the high Nd/Sr concentration ratios in the magma. Based on this hypothesis and on the reference age obtained from U–Th analyses of whole rocks and glasses, we suggest that the CI MT 1 resided some thousand years before erupting, and that the combined fractional crystallization/ assimilation process occurred during this time interval. The decreasing in Nd, Pb and (230Th/232Th) and the increasing is Sr isotope ratios (Figs. 6 and 7) shown by the CI MT 2 with respect to the pre-CI sample 947, suggests that contamination with continental crustal material with high Sr, and low Nd and Pb isotope and (230Th/ 232 Th) ratios, could have played a major role in the genesis of this CI end-member. We suggest that this process occurred at a crustal depths (~8 km), where poorly differentiated magmas were stored and differentiated to shoshonite-trachyte through combined fractional crystallization and assimilation of Hercynian crust. To constrain such process we used the EC-AFC (Energy Constrained Assimilation and Fractional Crystallization) model of Spera and Bohrson (2001). We assumed that a relatively poorly differentiated magma was stored at 8 km depth, where seismic data provide clear evidence for the occurrence of a low velocity layer that was interpreted as a partially molten zone (Zollo et al., 2008). We assigned to this magma a composition similar to the poorly evolved shoshonitic melts found in the CI melt inclusions with Sr, Nd, Th and Pb concentrations of 913, 48, 13.6 and 31.6 ppm, respectively (Webster et al., 2003; and unpublished own data). We further assumed Sr, Nd, and Pb isotope compositions based on the mean values of the data reported in D'Antonio et al. (2007) for the pre-CI magmas (Table C, Supplementary material). The (230Th/232Th) ratio of this magma is the age-corrected value measured for the pre-CI magmas dated at 47 ka which is enriched in 238U/230Th as a consequence of the source contamination process, as discussed above. As assimilant we used the average composition of the Hercynian continental crust of the Calabrian region (87Sr/86Sr = 0.71460, 143Nd/144Nd = 0.51222 and 206 Pb/204Pb = 18.46, Table C, Caggianelli et al., 1991; Rottura et al., 1991). The (230Th/232Th) was calculated at 0.78 based on secular equilibrium of typical crustal rocks ((230Th/232Th) from 0.6 to 0.8; Dickin, 1995; Fig. 6). This value is similar to the age-corrected (230Th/ 232 Th)ac values of 0.79 measured for the coarse-grained biotite

b

c

Fig. 7. a) Sr vs. Nd isotope ratios. Trend A models the Assimilation, Fractional Crystallization (AFC) process that produced, starting from a pre-CI magma, the least radiogenic and least differentiated CI MT 2 as a consequence of contamination with Hercynian crust; b) Trend B models the variations in terms of Th activity ratios and 143 Nd/144Nd isotope ratio due to the AFC process that produced the CI MT 2 as a consequence of contamination with Hercynian crust of a pre-CI magma; c) Trend C models the variations in terms of Pb and Sr isotope ratios due to the contamination with Hercynian crust of a pre-CI magma that produced the CI MT 2. The parameters used are in Table C, (Supplementary material).

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fraction (N1.5 mm) separated from the SMsp sample, and of 0.75 obtained for the Campi Flegrei more contaminated, least differentiated Minopoli 2 shoshonitic magma (B. Scheibner, personal communication). The equilibration temperature (Teq) was assumed equal to 950 °C, as this value is in the range defined by the results of thermometric experiments of pyroxene-hosted melt inclusions in CI products (Fulignati et al. 2004). At such temperature, the modeled daughter magma has Sr = 427 ppm, Nd = 46; Pb = 35 ppm, Th = 20 ppm, 87Sr/86Sr = 0.70731, 143Nd/144Nd = 0.51248, 206Pb/204Pb = 19.08 and (230Th/232Th) = 0.882. These values are close to those of the CI trachytic sample SMsp, representative of the MT 2 entering the shallow reservoir before eruption. The results of the modeling, obtained by using variable trace elements and isotope ratios, are reported in Fig. 7, and the parameters used in the simulation are listed in Table C. Fig. 7a shows that starting from a melt with Sr and Nd isotopes similar to the pre-CI samples, the isotope composition of the CI MT 2 can be obtained assuming that an EC-AFC process occurred at 8 km depth. The same process can explain the Th activity ratios (Fig. 7b) and Pb isotope ratios (Fig. 7c) of the least radiogenic CI MT 2. Based on both the previous discussion and modeling, we propose that the most evolved CI MT 1 was derived from a magma that differentiated mostly in a shallow storage region (Arienzo et al., 2009). By contrast, the parental melt of the CI MT 2 was stored and contaminated at a deeper crustal level within the Hercynian crust. After such contamination, this melt rose to shallower depth and crystallized at 45.4 ± 2.1 ka. This magma then entered a pre-existing shallow reservoir and mixed with the genetically and isotopically distinct resident magma (CI MT 1). Mixing likely occurred at 6.4± 2.1 ka preceding the large caldera-forming eruption at 39.3 ka.

composition as well as chemical and isotopic differences and medium to large scale compositional heterogeneity in the deposit. This short time window allowed the CI trachytic MT 2 to preserve its time information. The absolute age constraints, provided by the internal isochrons of the MT 2 suggest that the CI magma chamber and its large volume of chemically and isotopically distinct magmas, may not have been a long-lived system. Rather, the two CI end-member magmas assembled in a relatively short time before eruption, that is between 45.4 ± 2.1 and 39 ka. In this time interval crystallization, differentiation and mixing occurred resulting in the most cataclysmic calderaforming eruption in the Mediterranean area over the past 200 ka. Supplementary data associated with this article can be found, in the online version, at doi:10.1016/j.epsl.2011.04.002.

7. Conclusions

Acocella, V., Funiciello, R., Marotta, E., Orsi, G., de Vita, S., 2004. The role of extensional structures on experimental calderas and resurgence. J. Volcanol. Geotherm. Res. 129, 199–217. Allegre, C.J., Condomines, M., 1976. Fine chronology of volcanic processes using 238 U–230Th. Earth Planet. Sci. Lett. 28, 395–406. Arienzo, I., Civetta, L., Heumann, A., Wörner, G., Orsi, G., 2009. Isotopic evidence for open system processes within the Campanian Ignimbrite (Campi Flegrei—Italy) magma chamber. Bull. Volcanol. 71 (Issue 3), 285–300. doi:10.1007/s00445-0080223-0. Arienzo, I., Moretti, R., Civetta, L., Orsi, G., Papale, P., 2010. The feeding system of Agnano-Monte Spina eruption (Campi Flegrei, Italy): Dragging the past into the present activity and future scenarios. Chem. Geol. 270 (1–4), 135–147. Avanzinelli, R., Elliot, Tim, Tommasini, S., Conticelli, S., 2008. Constraints on the genesis of potassium-rich Italian volcanic rocks from U/Th disequilibrium. J. Petrol. 49 (2), 195–223. doi:10.1093/petrology/egm076. Bohrson, W.A., Spera, F.J., Fowler, S.J., Belkin, H.E., De Vivo, B., Rolandi, G., 2006. Petrogenesis of the Campanian Ignimbrite: implications for crystal–melt separation and open-system processes from major and trace elements and Th isotopic data. In: De Vivo, B. (Ed.), Volcanism in the Campania Plain: Vesuvius, Campi Flegrei and Ignimbrites. Dev. Volcanol. 9, 249–288. Bourdon, B., Zindler, A., Wörner, G., 1994. Evolution of the Laacher See magma chamber: evidence from SIMS and TIMS measurements of U–Th disequilibria in minerals and glasses. Earth Planet. Sci. Lett. 126, 75–90. Caggianelli, A., Del Moro, A., Paglionico, G., Pinarelli, L., Rottura, A., 1991. Lower crustal granite genesis connected with chemical fractionation in the continental crust of Calabria (Southern Italy). Eur. J. Mineral. 3, 159–180. Cannatelli, C., Lima, A., Bodnar, R.J., De Vivo, B., Webster, J.D., Fedele, L., 2007. Geochemistry of melt inclusions from the Fondo Riccio and Minopoli 1 eruptions at Campi Flegrei (Italy). Chem. Geol. 237 (3–4), 418–432. Cassignol, C., Gillot, P.Y., 1982. Range and effectiveness of unspiked potassium-argon dating: experimental groundwork and applications. In: Odin, G.S. (Ed.), Numerical Dating in Stratigraphy. Wiley, Chichester, pp. 159–179. Civetta, L., Orsi, G., Pappalardo, L., Fisher, R.V., Heiken, G., Ort, M., 1997. Geochemical zoning mingling eruptive dynamics and depositional processes — the Campanian Ignimbrite Campi Flegrei caldera, Italy. J. Volcanol. Geotherm. Res. 75, 183–219. Condomines, M., Sigmarsoon, P.J.G., 2003. Timescales of magma chamber processes and dating young volcanic rocks. Rev. Mineral. Geochem. 52, 125–174. Condomines, M., Hedmond, C., Allègre, C.J., 1988. U–Th–Ra radioactive disequilibria and magmatic processes. Earth Planet. Sci. Lett. 90, 243–262. Cooper, K.M., Reid, M.R., 2008. Uranium-series crystal ages. Rev. Mineral. Geochem. 69, 479–544. Costa, A., Dell’Erba, F., Di Vito, M., Isaia, R., Macedonio, G., Orsi, G., Pfeiffer, T., 2009. Tephra fallout hazard assessment at the Campi Flegrei caldera (Italy). Bull. Volcanol. 71, 259–273. D'Antonio, M., Tonarini, S., Arienzo, I., Civetta, L., Di Renzo, V., 2007. Components and processes in the magma genesis of the Phlegrean Volcanic District, southern Italy. In: Beccaluva, L., Bianchini, G., Wilson, M. (Eds.), Cenozoic Volcanism in the

U-series data presented here provide (a) absolute age information on a crystallization/differentiation event involving the least differentiated and least radiogenic CI MT 2, hence providing a minimum residence time for this melt, and (b) allow us to calculate a reference age for the most radiogenic and differentiated CI MT 1. Furthermore, based on Sr, Nd, and Pb isotope compositions in combination with (230Th/232Th) ratios measured on samples from pre-CI, CI and post-CI eruptions, we suggest that the two magmatic components that mixed during the CI eruption, evolved in different reservoirs through distinct differentiation processes and at different times. From the enriched source, magmas rose to crustal level and were stored in a deep reservoir (~8 km). From this reservoir, they rose to shallower depth (~5 km), where they were stored and continued to differentiate, partially preserving their Nd and Pb isotopic compositions, and where they eventually fed the pre-CI volcanism. We concur with Pabst et al. (2008) in suggesting that the parental magma of the CI MT 1 originated from one of these magmas. Magmas stored for longer time within the deep reservoir, differentiated and contaminated with the Hercynian crust, thus strongly modifying their Sr, Nd, and Pb isotope compositions and (230Th/232Th). In this reservoir the parental magma of CI MT 2 acquired its geochemical and isotopic features. When the regional and/or local stress regime allowed this deep contaminated magma to rise to shallower levels, it crystallized at around 45.4 ± 2.1 ka, forming large phenocrysts that are in Sr isotopic equilibrium with their host melt. This differentiation process produced a magma slightly more evolved than MT 2 and chemically similar to the SMc sample. MT 2, with its large amount of crystals, recharged a pre-existing shallow reservoir and partially mixed with the chemically most differentiated resident MT 1. Minerals exchanged mostly from the high temperature crystal-rich magma to the cooler crystal-free trachy-phonolitic MT 1. Mixing occurred over a relatively small time span and did not produced a complete homogenization of the CI magmas before eruption, as suggested by bimodality in glass

Acknowledgements The authors warmly thank G. Mengel for her technical support in the clean laboratory, A. Carandente and P. Belviso for their support during samples preparation. The authors thank S. de Vita for his support during field work, M. D'Antonio, B. Scheibner and S. Pabst for discussion and suggestion during this work. We are grateful to the Editor T.M. Harrison and two anonymous reviewers for their useful criticism and suggestions. We gratefully acknowledge the support by the Abt. Isotopengeochemie at GZG for providing access to the TIMS machine. This study benefited from the financial support of the 5th Framework Programme of the European Union within the ERUPT project.

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