Journal of Petrology Advance Access originally published online on March 31, 2006
Journal of Petrology 2006 47(8):1467-1507; doi:10.1093/petrology/egl015
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Evidence of Magma Mixing in the ‘Daly Gap’ of Alkaline Suites: a Case Study from the Enclaves of Pantelleria (Italy)
DIPARTIMENTO DI CHIMICA E FISICA DELLA TERRA ED APPLICAZIONI ALLE GEORISORSE E AI RISCHI NATURALI (CFTA), UNIVERSITÀ DEGLI STUDI DI PALERMO VIA ARCHIRAFI N, 36-90123 PALERMO, ITALY
RECEIVED APRIL 24, 2004; ACCEPTED MARCH 7, 2006
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONPETROGRAPHIC DESCRIPTIONSDISCUSSION OF MAGMA EVOLUTIONAPPENDIX: SAMPLE LOCATIONSREFERENCES |
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The island of Pantelleria consists of trachytes, pantellerites and minor mildly alkaline basalts. Rocks of intermediate composition (falling in the so-called ‘Daly Gap’) such as mugearites, benmoreites and mafic trachytes occur only in the form of enclaves in trachytes and pantellerites inside the main caldera of the island (Caldera ‘Cinque Denti’), which collapsed during the ‘Green Tuff’ ignimbrite eruption at
50 ka. The enclaves include volcanic, subvolcanic and intrusive rock types. The enclaves in host trachyte contain traces of glass; devitrified glass occurs within enclaves in host pantellerites. Minerals in the enclaves show regular compositional variations with whole-rock silica content. Glass present in the medium-grained samples is interpreted to be the result of incipient melting. The major and trace element compositions of the enclaves show regular and linear variations between an evolved mafic magma (hawaiite) and a felsic end-member similar to the ‘Green Tuff’ trachyte. Fractional crystallization modelling of compatible and incompatible trace elements (V, Ni, Zr, La, Sm, Lu, Nb, Y, Th) does not reproduce the observed trends. Rocks of intermediate composition within the ‘Daly Gap’ can be explained only by magma mixing between an already differentiated mafic magma (hawaiite) and an anorthoclase-rich trachytic melt in the lower and higher parts, respectively, of a stratified magmatic chamber. Medium-grained enclaves are interpreted as the result of fragmentation of solidified mixing layers in the roof of the magma chamber during the eruption of the ‘Green Tuff’, when the collapse of the caldera took place. Diffusion calculations suggest a residence time of <5 days for the enclaves in their host magmas. KEY WORDS: Daly Gap; enclaves; magma mixing; Pantelleria
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONPETROGRAPHIC DESCRIPTIONSDISCUSSION OF MAGMA EVOLUTIONAPPENDIX: SAMPLE LOCATIONSREFERENCES |
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Pantelleria is a Pleistocene volcanic island located in the Sicily Channel Rift Zone (Fig. 1) in which transtensional tectonics along the northern margin of the African plate have been active since the middle Pliocene. The island is constructed on crust constituting the foreland of the northern African plate, consisting of a deformed cover sequence (Tertiary–Pleistocene) overlying a Mesozoic carbonate succession and a thinned and unknown continental basement (Colombi et al., 1973
; Reuther & Eisbacher, 1985; Catalano et al., 1993; Berrino & Capuano, 1995). The subaerial part of the island consists of trachytes, pantellerites and minor alkaline to transitional basalts (Villari, 1974; Cornette et al., 1983; Civetta et al., 1984, 1988; Mahood & Hildreth, 1986).
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As frequently observed in many mildly alkaline suites in continental and oceanic areas, the erupted products lack rocks of intermediate composition, a phenomenon known as the ‘Daly Gap’ (Chayes, 1963
, 1977; Yoder, 1973). A number of non-conclusive petrogenetic models have been proposed to explain the Daly Gap (e.g. Baker, 1968; Cann, 1968; Clague, 1978; Lowenstern & Mahood, 1991; Bonnefoi et al., 1995; Mungall & Martin, 1995; Bohrson & Reid, 1997; Peccerillo et al., 2003; Avanzinelli et al., 2004). Important petrogenetic problems in the context of Pantelleria include: explaining the origin of the peralkaline silicic rocks, their relationship with the basaltic magmas, and the reasons for the scarcity of intermediate compositions in the volcanic sequence.
The study of magmatic enclaves from a suite of lavas characterized by a ‘Daly Gap’ can help to reveal the process responsible for the genesis of these particular rock fragments and can contribute to an understanding of the lack of intermediate erupted magmas at Pantelleria. The island is a favourable locality for this purpose because the erupted products are dominated by mafic (basalts–trachybasalts or hawaiites) and evolved (trachytes–pantellerites) rocks containing many enclaves of different composition.
In 1991, field research forming part of a geothermal project was carried out at Pantelleria during which magmatic enclaves were systematically collected. In this paper, we report on the petrological and geochemical study of these magmatic enclaves. Compositional data for whole-rocks and minerals are discussed with the aim of placing constraints on the genesis of the enclaves and explaining the transition between the mafic and felsic groups of magmatic rocks.
Geological setting
The island of Pantelleria consists mainly of lavas and welded ignimbrites of trachyte to pantellerite composition with subordinate basalts, and is the type locality for pantellerite and peralkaline volcanic rocks (Washington, 1913–1914). Large magnetic anomalies around the island suggest the presence of abundant basalt or gabbro intrusions at depth (Mahood & Hildreth, 1983; Berrino & Capuano, 1995).
The oldest exposed rocks are pantellerite lavas and ignimbrites with ages up to 324 ka. These are truncated by the La Vecchia caldera, which is dated at about 114 ka (Mahood & Hildreth, 1986). Post-caldera activity includes lavas, ignimbrites and pumice falls. Eruption of the ‘Green Tuff’, dated at 45 ka, led to the formation of the younger Cinque Denti caldera (Mahood & Hildreth, 1986), which, however, according to Civetta et al. (1988, 1998), is coincident with the northern part of the older La Vecchia caldera and the younger Monastero caldera (Fig. 1). The ‘Green Tuff’ is a unit compositionally zoned from pantellerite to trachyte and considered by Mahood & Hildreth (1986) and Civetta et al. (1988, 1998) as the result of a large eruption from a stratified magmatic reservoir. It is a major marker horizon representative of the first of six silicic eruptive cycles (Civetta et al., 1988, 1998), which lasted until 5·5 ka. Volcanic cycles I–VI are indicated in a schematic stratigraphic section in Fig. 1d. Subsequent eruptions filled the pre-existing caldera through several eruptive vents, the principal of which was at Monte Gibele; this volcanic edifice was subsequently tectonically uplifted and tilted to form Montagna Grande, the highest summit of the island (Washington, 1913–1914; Mahood & Hildreth, 1986). On the basis of K–Ar geochronological data, Mahood & Hildreth (1986) considered the Monte Gibele–Montagna Grande volcanic activity as a continuation of the ‘Green Tuff’ eruptive cycle; however, according to Civetta et al. (1988) the Montagna Grande is the result of eruptive cycle II. Considering the Pantelleria post-Green Tuff activity, mildly alkali basalts erupted sporadically (29 ka and <9 ka) solely from outside the calderas.
Sample collection and analytical methods
Samples of enclaves were grouped according to host-rock age based on the sequence of magmatic cycles proposed by Civetta et al. (1988). Enclaves were found and sampled almost exclusively from lavas that erupted in the area of the Cinque Denti caldera.
From 133 collected samples, often of small dimensions, 38 enclaves (x) and 22 host-rocks (h) were selected for detailed study on the basis of the following criteria (for the enclaves): greater size, lack of clay mineral alteration, negligible macroscopic interaction with the host-magma, hand specimen characteristics (colour, grain size) and the age of host-rocks (Tables 1 and 2). Among the widespread felsic enclaves <18 ka in age (Mahood & Hildreth, 1986; De Vivo et al., 1993), only a few representative samples were considered.
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The absence of secondary alteration in the enclaves was verified by X-ray diffraction (XRD) analysis using a Philips 1200 diffractometer at the Dipartimento CFTA of the University of Palermo (Italy); <2 µm oriented slides of water-centrifuged ground powders were analysed after natural, ethylene glycol, 300°C and 600°C heating treatments, respectively. For small enclaves, the host-rock was also analysed.
Modal proportions of minerals in the enclave and in the representative host-rock were determined by point counting involving 1631–2054 points covering the entire area of a thin section at 0·1 mm spacing with a 2–3% precision.
Mineral compositions were determined by electron microprobe analysis (EMPA) on an ARL-SEMQ at the Centro Studi Geominerari e Mineralurgici, CNR of Cagliari (Italy), with instrument settings at 24 kV accelerating voltage and 23 nA beam current with a diameter of 1 µm and counting times of 10 s. Brass and ZAF corrections were made using the MAGIC IV program (Colby, 1967) with precision better than 5%. Glass and feldspar were analysed with a 10 µm defocused beam on a vibrated sample to avoid analytical bias on the Na determination by EMPA (Spray & Rae, 1995; Olmi, 1997). For Cl, Sr, Ba, and F determination the detection limit was >30–40 ppm.
Whole-rock analyses were carried out on samples initially washed in distilled water to minimize the effect of marine aerosol. Major and trace elements were determined by X-ray fluorescence (XRF) at the Dipartimento CFTA of the University of Palermo, on pressed pellets using a wavelength-dispersive XRF Philips PW 1400 spectrometer with Cr and Rh tubes. A matrix effect correction procedure was performed using the experimentally determined coefficients (Franzini et al., 1975; Leoni & Saitta, 1976; Leoni et al., 1982). Nb was measured by a step-by-step slow scan. The uncertainty for major elements is <1% for Si, Al and Na, <3% for Ti, K, Fe and Ca, and <10% for Mg, Mn and P. The uncertainty for trace elements is estimated to be <7% for Rb, Nb, Y, Sr and V, and <15% for Zr, La, Ce, Ba, Ni, Cr and Cl.
Several trace elements for some II Cycle enclaves, selected according to the available amount of sample, were determined by inductively coupled plasma mass spectrometry (ICP-MS) at the Dipartimento ITAF of the University of Palermo, using a Hewlett Packard 4500. The precision is 2–10% for trace element contents in the range 10–150 ppm and 5–30% for trace element contents in the range 0·1–5 ppm. Loss on ignition (LOI) was determined by gravimetric methods. mg-number = Mg/(Mg + Fe2+) and Fe3+/Fe2+ ratio were calculated according to the method of Kilinc et al. (1983).
Volcanic enclaves were classified according to the total alkali–silica (TAS; Le Maitre et al., 1989) and Macdonald (1974) schemes; medium-grained enclaves were classified according to De la Roche et al. (1980) and MacKenzie et al. (1982) diagrams.
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PETROGRAPHIC DESCRIPTIONS |
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TOPABSTRACTINTRODUCTIONPETROGRAPHIC DESCRIPTIONSDISCUSSION OF MAGMA EVOLUTIONAPPENDIX: SAMPLE LOCATIONSREFERENCES |
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Size and shape
The enclaves vary in size from centimetres to decimetres and show irregular, rounded or ellipsoidal shapes. There appears to be no relation between the shape of the enclaves and their internal structure, suggesting that they were solid fragments when incorporated in the host magma. Enclaves hosted in trachyte and pantellerite seem to have different shapes and sizes; the former exhibit margins grading into the host-rock and are frequently rounded, whereas the latter display sharp edges and have irregular shapes (Fig. 2).
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Texture
Textures in the enclaves vary from holocrystalline, fine- to medium-grained, to porphyritic; some contain traces of glass. Enclaves have been grouped according to their texture, into volcanic, fine- to medium-grained and medium- to coarse-grained types. As noted above, there is a distinction between enclaves in pantellerite hosts and enclaves in trachyte hosts. Within the glass-bearing types it is possible to distinguish those with glass with abundant microlites, glass with few microlites, and clear to dark glass without microlites.
Modal mineralogy of the enclaves
The results of the modal analysis for selected enclaves within the trachytes and pantellerites are presented in Table 1. Enclaves within trachytes have the following mineralogy: (a) mafic enclaves: Ol (10–3), Cpx (33–13), Feldspar (65–33), Fe–Ti oxides (5–3), Apatite (tr; trace), glass (17–9); (b) intermediate composition enclaves: Ol (8–2), Cpx (25–6), Feldspar (79–49), Fe–Ti oxides (9–3), Apatite (tr), glass (36–tr); (c) felsic enclaves: (c1) trachyte enclaves: Cpx (2–0), Anorthoclase (20–17), groundmass (81–78); (c2) microgranite enclaves: Feldspar (60–47), Qz (35–22), Bi (tr), Fe–Ti oxides (tr), glass (tr), quartz and feldspars are normative; syenite and pantellerite enclaves were not found. Enclaves within pantellerites are as follows: (a) trachyte enclaves: Cpx (2–1), Anorthoclase (26–18), Fe–Ti oxides (tr), Ol (tr), groundmass (72–79); (b) syenite enclaves: Ol (2–0), Cpx (7–tr), Anorthoclase (88–17), Amph (3–0), Bi (tr), Fe–Ti oxides (tr), groundmass (81–0); (c) pantellerite enclaves (one sample) Anorthoclase (20), Cpx (1), Fe–Ti oxides (tr), Cossyrite (tr), glass (79); (d) microgranite enclaves (one sample) Alkali feldspar (63), Qz (26), Bi (tr), Fe–Ti oxides (tr), quartz and feldspars are normative. Enclaves of intermediate composition in host pantellerite were sporadically found in the intra-caldera ignimbritic pantellerite, e.g. sample 31x, which is very similar to the intermediate enclaves in the trachytes. We emphasize, in this context, the wide glass content variation from 0% up to 36% in the medium-grained sample 25x. Glass is generally more abundant towards the margin of the enclaves. The result of modal analyses for representative enclave host-rocks are presented in Table 2.
Enclave types
Volcanic enclaves
These include rock types similar to the common volcanic rock types of Pantelleria, including both trachytic and pantelleritic compositions.
Fine- to medium-grained enclaves
Although traces of interstitial glass occur in some fine- to medium-grained enclaves, the majority have textures similar to those of sub-volcanic rocks.
The more felsic types have textures ranging from granophyric to porphyritic and show some similarity to trachytes and pantellerites, except for their crystalline groundmass. They contain phenocrysts of zoned anorthoclase, subordinate Na–Fe-hedenbergitic clinopyroxene, fayalitic olivine and cossyrite in a microcrystalline groundmass of Na-sanidine, quartz, cossyrite, green aegirine-rich clinopyroxene, opaque minerals and apatite.
The more mafic enclaves have textures ranging from doleritic to porphyritic with phenocrysts of equant, forsteritic, olivine in an sub-ophitic groundmass of labradorite–andesine plagioclase, Ti-augite, Ti-magnetite, ilmenite, apatite and interstitial dark glass (Fig. 3a). Enclaves of hawaiitic composition contain spongy centimetre-sized xenocrysts of anorthoclase (Fig. 3a) and are found in volcanic rocks >50 ka in age, from the II, IV and VI cycles. These enclaves are similar, apart from the crystallization state, to the mafic sample 33MB from C.da Mursia of Mahood & Baker (1986) considered ‘a result of basalt contamination by trachyte or syenite or of a mixing process with a trachytic magma’.
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Medium- to coarse-grained enclaves
These rocks include alkaline granites, syenites, syenodiorites and syenogabbros with crystals of millimetre size. They resemble intrusive rocks but sometimes have interstitial glass.
The alkaline granites have porphyritic textures with zoned anorthoclase, hedenbergitic pyroxene with an aegirine-rich rim, quartz, alkali amphibole and biotite.
The syenites are common in the host trachytes of the earliest pre-Green Tuff volcanic eruptions, as well as the I and IV cycles. They consist of zoned anorthoclase with a resorbed core and green, zoned, hedenbergitic clinopyroxene, fayalitic olivine, interstitial Fe–Ti oxides, apatite, aegirine–augite, green to reddish brown pleochroic amphibole, biotite and quartz.
The syenodiorites show textural similarities to orthocumulates (Fig. 3d). Such enclaves, found in the II and VI cycle rocks from the Monastero caldera, were previously described as ‘benmoreites’ (Villari, 1974; Civetta et al. 1984). They show different mineral generations (e.g. relics of equant olivine with opaque inclusions and zoned elongate to skeletal olivine). The clinopyroxenes occur either as lobate augite with opaque inclusions and Ti-rich rims, or as clear, Fe-rich, zoned pyroxenes with augitic cores. Green, aegirine-rich rims are locally observed. Most feldspar is zoned with a clear, twinned, plagioclase core and a rim of zoned to xenomorphic alkali-feldspar. Ti-magnetite, ilmenite and needles of apatite are also present.
Syenogabbro and monzonitic gabbro enclaves are rare in the oldest pantellerites. They are similar to the syeno-diorites with a comparable intercumulus assemblage of pleochroic alkaline amphibole, acicular aegirine-rich clinopyroxene and biotite.
Glass–enclave relations
The enclaves contain a variety of types of glass (Figs 3 and 4):
- clear to dark glass with microlites, occurring as localized areas in the outermost parts of the enclaves, suggests leakage of the host magma into the enclave;
- brown glass with few microlites, often exhibiting feathery crystallization, quench skeletal or radial crystals, locally with bifid edges, is distributed along grain boundaries in the enclaves; this kind of glass is often found close to mafic minerals;
- clear glass in contact with brown to dark glass occurs in the interiors of the enclaves, without microlites.
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Enclave–host magma relations
The granular enclaves within pantellerite lavas (e.g. the syenite sample with amphibole 81x, the microgranite 79x, and the syenodiorite enclave 31x) do not contain glass. In the fine-grained types, portions of the mesostasis in trachyte and pantellerite enclaves show traces of rapid devitrification of the glass into globular or fibrous clusters of Na-sanidine, quartz, aegirine-rich clinopyroxene and cossyrite. Fibrous feldspars nucleate around the rims of anorthoclase phenocrysts (Fig. 3 h), which are often broken along cleavage planes or fragmented into small subidiomorphic quadrate grains. In holocrystalline medium-grained and hypidiomorphic types, alkali feldspar shows antiperthite exsolution.
Enclaves in trachytes are commonly characterized by the presence of interstitial glass and deformations of crystals, but show distinctive differences according to the enclave type: (1) The enclaves with a bulk-rock composition of pantellerite (only in pre-Green Tuff trachytes) contain glass varying from colourless to dark brown, in some cases, similar to the glass derived from the host magma; (2) in the mafic doleritic enclaves, glass develops as inclusions in spongy anorthoclase xenocrysts and interstitially in the groundmass; (3) the medium- to coarse-grained enclaves exhibit glass very similar to the residual intra-crystalline melt, changing from colourless to darkish brown. The sharp colour change occurs on a <10 µm interval, which is indicative of very low interdiffusion of elements between the two different glasses (Figs 3 and 4). These enclaves are poly-crystalline aggregates with allotriomorphic-granular textures in which radial intergrowths (MacKenzie et al., 1982) of pyroxenes and feldspars are occasionally observed; however, the side of these feldspar crystals in contact with the surrounding glass is idiomorphic with sharp boundaries (Fig. 3 g). Feldspar crystals in contact with the glass show evidence of rim resorption and euhedral, more anorthitic, cores (Figs 3d,e and 4). Both crystals and interstitial glassy areas are crossed by needles of skeletal olivine, pyroxene and even apatite (Fig. 3e and f). According to Ferla et al. (1978) and Harris & Bell (1982), all these textures are consistent with heating and partial remelting of the enclaves by the host magma. Glass in the enclaves is nearly absent when the host trachyte is holocrystalline, as in some Montagna Grande trachytic rocks.
Bulk-rock geochemistry of the enclaves
Bulk-rock chemical analyses of the enclaves and their host lavas are reported in Tables 3–5 and compared with available published data for Pantelleria rocks in Fig. 5a–c (Carmichael, 1962, 1967; Rittmann, 1967; Romano, 1969; Macdonald, & Bailey, 1973; Villari, 1974; Civetta et al., 1984, 1988, 1998; Mahood & Baker, 1986; Mahood & Stimac, 1990; Esperança & Crisci, 1995; Avanzinelli et al., 2004). According to the TAS classification diagram (Le Maitre et al., 1989) the enclaves span the compositional range between the mafic and felsic volcanic rocks, corresponding to the ‘Daly Gap’, typically found between basalts and trachytes (Fig. 6). Volcanic or medium- to coarse-grained enclaves of trachytic and pantelleritic composition are also present.
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R2 <0·3) mass-balance calculations for major oxides (
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The volcanic rocks of Pantelleria typically show a bimodal distribution of compositions with hawaiites and trachytes as mafic and felsic end-members, respectively, and a wide ‘Daly Gap’ between 52% and 63% SiO2. The host-rocks to the enclaves are trachytes with at least 61% SiO2, almost exclusively of the II cycle, and subordinate pantellerites. Enclaves of basaltic composition or mafic cumulates have not been found.
Element variations with silica content and rock type
Major and trace element Harker variation diagrams for enclaves with SiO2 ranging from 48% to 65% display, with few exceptions, distinctive linear and regular trends that are different from those of the erupted volcanic rocks (Fig. 5). Among the enclaves and volcanic rocks of Pantelleria we have, therefore, distinguished three chemical groups based on SiO2 content: mafic (45–48 wt % SiO2), intermediate (48–65 wt % SiO2) and felsic (65–73 wt % SiO2) (Figs 5 and 6). A plot of the agpaitic index (AI) [molecular (Na2O + K2O)/Al2O3) vs SiO2 (whole-rock per cent)] for the enclaves (Fig. 5b) shows a linear trend for AI values <1, varying from AI 
0·5 for mafic enclaves to AI 1·1 for enclaves with 65 wt % SiO2. The felsic volcanic rocks show a different trend, with progressively increasing peralkalinity (AI >1) with increasing silica content. The majority of the enclaves fall within the intermediate composition range, whereas the majority of the volcanic rocks are felsic (comenditic trachytes and pantellerites). We emphasize that the enclaves of intermediate composition (intermediate enclaves) practically are hosted solely in the II cycle trachytic volcanic rocks; however, the felsic end-member of the intermediate group is a syenite (81x) within a pantellerite host-rock.
Mineral chemistry and mineral–melt equilibria in the enclaves
The chemistry of minerals from the studied enclaves was compared with that for the volcanic rocks of Pantelleria available from the literature.
Olivine
Olivines with the highest Fo (Table 6, Fig. 7) contents correspond to the cores of crystals or to euhedral olivines coexisting with skeletal olivines. The skeletal crystals have lower Fo contents. Excluding the relict equant olivine (Fo86–84), the rim or skeletal olivines of the last generations to crystallize vary in composition with increase in whole-rock silica from Fo79–80 and 49·22 wt % whole-rock SiO2 (21x) to Fo2–6 in syenite (81x).
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The Mg-rich cores of olivine in enclave 21x (Fo86) (ol1 in Fig. 3a) are compatible with crystallization from a fairly primitive melt with mg-number
0·65 (Snyder & Carmichael, 1992) and a calculated temperature of crystallization, on the basis of the olivine–liquid geothermometer, of 1230°C [applying the method of Roeder & Emslie (1970)], or of 1215°C [according to Sisson & Grove (1993b)], under dry conditions and low pressure. In comparison, the basalts of Pantelleria exhibit a mg-number range from 0·58–0·54 (Mahood & Baker, 1986). The calculated liquid (FM-1) in equilibrium with the (Fo79) olivine rim from sample 21x (ol2 in Fig. 3a; Table 6), which contains anorthoclase, has mg-number 0·56 and a temperature of 1193°C, according to the liquid geothermometer of Sisson & Grove (1993b). This value is similar to that for the less-evolved basalt 38 of Mahood & Baker (1986), for which a liquidus temperature of 1181 ± 5°C was experimentally calculated at low pressure: the reliability of the Sisson & Grove (1993b) liquid geothermometer was also tested on basalt 38MB composition, for which the same temperature value of 1181°C has been calculated.
The estimated temperature of crystallization for the intermediate composition unzoned olivines (ol3 in Figs 3b and 7, and Table 6), calculated according to the olivine–liquid geothermometer of Sisson & Grove (1993b), varies from 1082° to 948°C with increasing silica content of the host-enclave; these olivines haveFe/MgKDol-w.r. >0·3, in agreement with the Carmichael & Ghiorso (1990) calculations for olivine with Fo <58. Temperature values close to 855 ± 10°C were calculated for the skeletal olivine and the adjacent surrounding glass, and can account for a not-approached equilibrium or for the probable limit of the olivine–liquid geothermometer, in agreement with that reported by Sisson & Grove (1993b).
The calculated liquid (FM-1) can be obtained by mass-balance calculations subtracting a 5 (wt %) of Fo79 olivine (rim) and a further 20 (wt %) of trachyte composition, from the 21x whole-rock. The primitive melt composition (FM-0) with mg-number 0·65 can tentatively be obtained by mass-balance calculations adding a 10 (wt %) of Fo86 olivine (core) to the FM-1 calculated liquid (Table 6). As a consequence of these calculations it is possible to assert that the 21x doleritic enclave, and also the 16x enclave, display traces of olivine xenocrysts of a more mafic basaltic parental magma.
Clinopyroxenes
Different kinds of clinopyroxene crystals are present in the enclaves. Some are partially resorbed, subhedral, equant crystals containing opaque micro-inclusions; more frequently, they show prismatic, elongated shapes, which can coexist with equant types. Sometimes skeletal clinopyroxenes occur in radial intergrowths with plagioclase (Fig. 3 g). Similar to the olivine trend (Fig. 7), except for the relict crystals, the rims or skeletal pyroxene of the last generations to crystallize vary in composition with an increase in whole-rock silica (Table 7, Fig. 8). Compositions range from Ca-rich clinopyroxene in dolerite to ferro-augite to Na-ferrohedenbergite in intermediate enclaves and aegirine–augite in syenite and felsic volcanic rocks. An increase of Fe/Mg and MnO (wt %) and a decrease of Al2O3 (wt %) and TiO2 (wt %) was observed in clinopyroxenes with increasing whole-rock silica content (Fig. 8). The doleritic enclaves contain sub-ophitic pyroxenes with high Al and Ti contents; these are typically indicative of high rates of crystallization (Feeley & Dungan, 1996).
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0·7 (
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Clinopyroxenes with mg-number >0·7 [‘basaltic’ according to Nimis (1995)
], are present in the more mafic enclaves, both as cores of zoned crystals and as xenocrysts coexisting with pyroxenes that have a higher Fe/Mg ratio. In many cases we have noted an inverse zoning pattern with a sudden increase of Al and Ti followed by an equally strong decrease of the same elements towards the rims of the pyroxenes. If we consider Fe/MgKDcpx-liq = 0·22 ± 0·04, according to Sisson & Grove (1993a, 1993b), the Mg-rich cores of some Cpx in intermediate enclaves (Table 7) with mg-number >0·7 (e.g. 16B cpx), are not in equilibrium with the whole-rock (liquid). These xenocrysts, applying the geothermometer of Putirka et al. (1996), are in equilibrium (KD = 0·26) with a calculated basaltic liquid FM-1 (Table 6) at about 1168 ± 2°C and <0·5 GPa pressure. A similar temperature was determined experimentally by Mahood & Baker (1986) for cpx crystallization from a Pantelleria basalt at low pressure. Applying the clinopyroxene geobarometer of Nimis (1995), pressures of <0·2 GPa were obtained for the cpx cores in syenogabbro and syenodioritic enclaves (Table 7).
The skeletal clinopyroxenes in the intermediate composition enclaves could be in equilibrium with a liquid represented by the whole-rock, at estimated temperatures varying from 1070° to 987°C with increasing whole-rock silica content, at low pressure, taking into account the uncertainty that derives from the estimated composition of the original liquid. The lowest equilibrium temperature range, for clinopyroxene in the syenodioritic enclaves, is 994–981°C. These values have been calculated for the Fe–Mn-rich rims of the zoned cpx and for the skeletal pyroxenes in radial intergrowths with the plagioclase (oligoclase), suggesting an equilibrium liquid with a composition similar to that of the intracrystalline glass (e.g. 22x). It must be noted that the estimated temperature ranges for intermediate clinopyroxenes and olivines give fairly comparable results, which are also confirmed by the olivine–augite geothermometer of Loucks (1996) with a 985 ± 6°C estimated temperature. The aegirine–augite clinopyroxene appears as a late skeletal crystallization product in devitrified glass.
Feldspars
Feldspar compositions vary with-whole rock silica content (Table 8, Fig. 9); they are An58–50 in the dolerite enclaves. The plagioclases of the intermediate enclaves vary from andesine in the zoned euhedral cores to oligoclase or anorthoclase in the rims (e.g. An45–An19–An6 range in 20x feldspar). Alkali feldspar (Or38), close to the binary minimum (Carmichael & MacKenzie, 1963), is present in the syenitic enclaves on the rim of the anorthoclase and in the groundmass of the trachytic enclaves.
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All enclaves with intermediate compositions between basalt and trachyte contain resorbed anorthoclase with ovoid cores that we consider xenocrysts (Fig. 3c). The enclaves without glass contain anorthoclase in the rim overgrowths on plagioclase and resorbed megacrysts of earlier-formed anorthoclase.
It must be noted that the alkali feldspar in the trachyte-hosted enclaves does not display the perthite textures that are observed in the enclaves within the pantellerites. Trachytes contain An10–2Or20–28Ab65 anorthoclase and, in places, Na-sanidine in the groundmass. The Ba content in the feldspars of the enclaves increases while the An content progressively decreases (not shown). Feldspars show, in places, patchy-zoning: normally zoned feldspars (62x) coexist with crystals retaining rounded cores of ternary feldspar, euhedral external cores of An40 plagioclase and an outermost zone of clear alkali-feldspar. According to Sisson & Grove (1993a) and Feeley & Dungan (1996), the plagioclases from the intermediate enclaves could be in equilibrium with a liquid similar to the whole-rock composition, under almost dry conditions (KD = 1 – 2) and at low pressure (<0·2 GPa) (Fig. 9b). However, the Ca-rich composition of the hypothesized liquid is probably due to the presence of traces of Ca-rich clinopyroxene and Ca-rich plagioclase (>An70) xenocrysts. The andesine–labradorite plagioclases (A58–38), in places occurring as resorbed feldspar cores, do not contain apatite inclusions and have compositions very similar to the plagioclase synthesized in the 1050 ± 25°C range in the experiments of Mahood & Baker (1986). For the radial textural intergrowth of oligoclase plagioclase An20–18 and Fe–Mn-rich skeletal clinopyroxene the equilibrium temperature must be the same: the calculated temperature range for clinopyroxene is 994–981°C and is surprisingly very similar to the 990–960°C temperature range, calculated according to Mathez (1973), for oligoclase in equilibrium with a liquid with the intracrystalline glass composition (e.g. 22x) (Fig. 9c). The Ca-rich plagioclases are not in equilibrium with interstitial glass (Fig. 9c).
Fe–Ti oxides
Many researchers have reported the occurrence of coexisting Fe–Ti oxide pairs, ilmenite and magnetite, in all the rocks of the island. Only Fe–Ti oxide pairs with equilibrium Mn/Mg partitioning according to Bacon & Hirschmann (1988) were considered here (Table 9).
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In a plot of T°C vs fO2, calculated using the method of Ghiorso & Sack (1991)
(Fig. 10), the enclave and volcanic rock data plot subparallel to the quartz–fayalite–magnetite (QFM) buffer for temperatures up to 1050°C, corresponding to the equilibrium conditions of mafic melts. The data display a decrease of 
NNO (where NNO is the nickel–nickel oxide buffer), according to Carmichael (1991), from hawaiite to trachyte and trachypantellerite, and intercept the trend defined by the pantellerite data of Carmichael (1991) and Lowenstern et al. (1993), and the experimental data for samples from Ferdinandea Island (Graham Bank, Sicily Channel) reported by Carapezza et al. (1979). A temperature of 1075°C was determined by Mahood & Baker (1986) for an experimental differentiated mafic melt (Fliquid = 0·47, hawaiite); the highest temperature calculated for the intermediate composition enclaves is close to 1050°C (16x; mugearite). Temperatures of 1003–987°C were calculated from Fe–Ti oxide pairs associated with Fe-augite pyroxene, and 886–864°C from oxide pairs associated with aegirine–augite pyroxene. All the analysed enclaves containing interstitial glass, enclosed within vitrophyric lavas of trachytic composition, give temperatures between 1000 and 950°C.
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Amphiboles
Previous studies on the felsic volcanic rocks of Pantelleria revealed the presence of alkali amphiboles of barkevikite and arfvedsonite composition (Rittmann, 1967
; Villari, 1974). Ti–Na-arfvedsonite and Fe-richterite have been found in trachytic pantellerite and pantellerite rocks, respectively (Avanzinelli et al., 2004). Fe-richterite [classified according to Leake et al. (1997)] is also found in the comenditic trachytes (Civetta et al., 1998). In the syenite enclaves we have found xenomorphic Fe-richterite with Na increasing towards the rims of the zoned crystals (Table 8).
Apatite
Apatite, in the intermediate composition enclaves, is found enclosed solely in the external parts of zoned feldspars and surrounding glass, but it is absent in plagioclase cores. In doleritic enclaves it is confined within a matrix containing brown glass. In medium-grained enclaves, syenogabbros and syenodiorites, apatite crystals are needle-shaped with a morphological axial ratio of c:a >50:1 (Fig. 3e and f), whereas in the syenite enclaves the apatite crystals have a more stumpy shape. The high rate of crystallization is responsible for the morphology of the acicular apatite (Frost & Mahood, 1987). In this context, the presence of needle-shaped apatite crystals as the only microlites in the glass should be noted (Fig. 3e).
Cossyrite and quartz
The mineral cossyrite (aenigmatite), named after the pre-Byzantine name of Pantelleria, Cossyra, was detected in the pantellerite hosting the syenite enclaves and shows an almost constant composition comparable with data from the literature (Carmichael, 1962; Zies, 1966; Wolff & Wright, 1981; Mahood & Stimac, 1990; Civetta et al., 1998; Avanzinelli et al., 2004) (Table 9). Quartz was observed in the syenite enclaves and in the more felsic rocks (based on X-ray data).
Glass composition
The small deficiency in the sum of some of the analyses, considering the total Fe as Fe2O3, might be due to the presence of unanalysed Cl, Fl, Sr and Ba (Table 10).
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The glass in the doleritic enclave (21x), with hawaiitic composition, is peralkaline and silica-undersaturated, with a higher TiO2 content than the other analysed glasses. The 21x glass has the highest K2O content of all the known rocks of Pantelleria and a K2O/Na2O ratio of 0·75, which is very similar to that of the glass in the trachytic and pantelleritic volcanic rocks studied by Mahood & Stimac (1990)
. This glass, representing 17·5 vol. % of enclave 21x (Table 10), has a particularly high F and Cl content; the value of 9400 ppm Cl detected in the glass corresponds to 1645 ppm in the bulk-rock enclave. This provides evidence of the incompatible behaviour of Cl, which is almost totally concentrated in the melt. The total Cl content of the 21x hawaiitic enclave is comparable with that of other mafic rocks (see Fig. 5b). Volcanic rocks with compositions similar to the glass in 21x have not been discovered on the island.
The glasses in the intermediate and medium-grained enclaves, syenogabbros and syenodiorites are silica-saturated with compositions ranging from metaluminous, alkaline to peralkaline quartz-trachyte to pantellerite on an R1–R2 diagram (not shown) after De la Roche et al. (1980). The glasses appear similar in composition to the felsic volcanic rocks of Pantelleria, with AI = 0·94 ± 0·21. Considering the measured whole-rock Cl content as all concentrated in the glass, the calculated equivalent bulk-rock contents are in agreement, according to Webster (1997), with the chlorine solubility in felsic melts, in the temperature range 800–1075°C and at pressures of 0·1 MPa to 0·2 GPa. There are significant heterogeneites within the glass in individual enclaves (Table 10). However, different colours are often observed in the same glass area (Figs 3 and 4) and do not appear to be related to compositional variations. Mg-rich olivine, Ca-rich pyroxene and Ca-rich plagioclase are not in equilibrium with the surrounding glass. Traces of glass with a very high silica content (SiO2 74–76 wt %) were found in the syenogabbro (16Bx) enclave hosted in a trachyte. Syenite enclave (81x) within a hyalo-pantellerite host lava does not contain glass. We have found glass in the angular trachytic and pantelleritic wall-rock enclaves in host pyroclastic pantellerite. The glass within the spongy anorthoclase (22x) is similar to the external glass.
The behaviour of Sr in the Pantelleria magmas is strongly controlled by plagioclase crystallization, with a progressive lowering of Sr contents from basalts to trachytes, down to the very low concentration in the pantellerites. The Sr content of the analysed glass at the contact with the spongy anorthoclase enclaves (e.g. 22x) is variable and ranges from >1000 ppm down to below detection limits (<40 ppm); the latter corresponds to the low Sr content of the host lava.
Applying the liquid geothermometer of Sisson & Grove (1993b) to the analysed glass compositions and considering, for the original melt, H2O content of 0·5–1·85 wt %, the results are 938–983 ± 14°C and 876°C respectively for the trachytic and the more felsic glass compositions (e.g. the glass in 16Bx syenogabbro). The reliability of the Sisson & Grove (1993b) liquid geothermometer was also tested on pantellerite compositions: it must be noted in this context that the same temperature (850°C) was calculated for the 81 h and 32MS hyalopantellerites (Table 12), and for the 3112CG pantellerite, all with a pre-eruptive H2O content of 1·85 wt %, and that the same temperature value of 850°C was estimated from the Fe–Ti oxide geothermometer for 3112CG pantellerite (Carmichael & Ghiorso, 1990; Lowenstern et al., 1993) with a pre-eruptive H2O content of 1·4–2·1 wt % (Lowenstern & Mahood, 1991).
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DISCUSSION OF MAGMA EVOLUTION |
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TOPABSTRACTINTRODUCTIONPETROGRAPHIC DESCRIPTIONSDISCUSSION OF MAGMA EVOLUTIONAPPENDIX: SAMPLE LOCATIONSREFERENCES |
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One of the main problems in studies of peralkaline suites, such as those of Pantelleria, is the need for an explanation for the lack of rocks of intermediate composition, constituting the so-called ‘Daly Gap’. Among the several classes of hypotheses proposed to explain this distinctive feature, the most common are based on fractional crystallization models, with variants involving compositionally stratified magma chambers (e.g. Barberi et al., 1975
; Civetta et al., 1998; Peccerillo et al., 2003). Other hypotheses suggest that the peralkaline silicic rocks and associated basalts represent two different magma types that are not petrogenetically related (Mahood et al., 1990; Lowenstern & Mahood, 1991).
Fractional crystallization modelling of the enclaves spanning the ‘Daly Gap’
A step-by-step magmatic differentiation process from basalt to trachyte has been modelled, where at each stage the more progressively differentiated daughter magma represents the parental magma of the subsequent differentiation stage. The mineral compositions in the rocks 1MB basalt, 28x hawaiite, 16Bx mugearite, 22x benmoreite and 108x trachyte were used to model the phenocryst extract during progressive differentiation (Table 11). Using these mass-balance results, theoretical differentiation trends for both compatible and incompatible trace elements (Ni, V, Th, La, Sm, Lu, Nb, Y, Zr,) were calculated on the basis of Rayleigh fractionation for the complete step-by-step differentiation model covering the ‘Daly Gap’, and compared with the observed trends. Although a successful mass balance can be achieved for major element oxides, the trace element trends (Fig. 5c) cannot be reproduced.
We have also tested the fractional crystallization hypothesis by considering plots of an incompatible element vs the ratio of the same incompatible to a less-incompatible element (Joron & Treuil, 1989), such as the plots of Hf vs Hf/La, Zr vs Zr/Y, Th vs Th/Sm (Ce vs Ce/Yb, Cs vs Cs/Yb, Nb vs Nb/Y not shown) in Fig. 11. These diagrams illustrate that the enclaves, which range from mafic to felsic in composition, define a coherent, but much steeper, trend than the lavas.
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Magma mixing modelling: hawaiite and trachyte as end-members of intermediate composition magmas
A linear array of data in major element variation diagrams can be interpreted as the result of mixing between two magmas with different compositions (Bacon, 1986
; Frost & Mahood, 1987; Gourgaud & Villemant, 1992; Feeley & Dungan, 1996). In the Harker diagrams (Figs 5 and 12a) in the SiO2 48–65% range, the enclaves show at least two distinctly different trends of MgO, P2O5, Cr, Ni and Sc (not shown). In the majority of the enclaves the MgO contents define a negative trend with good linear correlation, which intercepts the trend defined by the mafic lavas at SiO2 48% and MgO 4% (hawaiite) and the field of felsic lavas at SiO2 65% and MgO <0·5% (trachyte) (Fig. 5a).
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The calculated mafic end-member (Table 12) is a hawaiite that nearly coincides with the bulk-rock analysis of an enclave (103x) from the >50 ka volcanic rocks with MgO 3·85% (Trend I, Fig. 12a), and comparable with the composition of the differentiated magma (F = 0·50, T = 1082°C and QFM conditions) calculated by Mahood & Baker (1986)
on the basis of an experimental study at low pressure of a less evolved basalt (38MB). Applying the Sisson & Grove (1993b) geothermometer to the calculated mafic end-member gives a similar temperature if we consider that the original melt had 0·5–1·5 wt % H2O. The felsic end-member corresponds to SiO2 65% and MgO < 0·5%, and a calculated temperature of 1000–950°C.
The doleritic enclaves, the basalts with anorthoclase xenocrysts, hawaiite 33MB and two intermediate composition enclaves SIC11C and SIC52C, define a second array in Fig. 12 that intercepts the mafic lava trend at SiO2 48% and MgO 7% (1190°C) and the felsic trend at SiO2 65% and MgO < 0·5%. This second trend (trend II in Fig. 12a) is probably formed by the mixing between an FM-1 basalt, probably rising along dykes, and overlying trachytic magma. The major and trace element contents of the mixing end-members, calculated by taking into consideration the 48% and 65% SiO2 intercepts of the least-squares calculated linear trends, or by linear extrapolations, are given in Table 12.
The hypothesized mixing model between hawaiite and trachyte (trend I) is consistent for all the analysed elements (for the REE plot of intermediate enclaves see Fig. 13) and is in perfect agreement with all the analysed enclaves of intermediate composition. Some differences between the measured and calculated data are probably due to analytical bias for low trace element contents, or to overlaps of the data with trend II.
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The origin of the glass in the enclaves
We can distinguish three different and simple hypotheses for the origin of the glass in the intermediate enclaves: (1) the enclave is partially crystallized after magma mixing, with the glass representing the chilled mixed melt, when it is incorporated in a partially molten/plastic state in the host magma; (2) the glass in the enclave derives from host magma infiltration; (3) the interstitial glass is the product of partial remelting of previously solid enclaves heated by the host magma. We will evaluate each of these hypotheses in turn.
(1) The glasses in the intermediate enclaves are compositionally very different from those of mixtures of hawaiitic and trachytic melts: they have trachytic or felsic compositions. The inhomogeneous melt is not comparable with an intercumulus liquid within a crystal mush according to the following considerations. (a) A syenogabbro enclave (16B), hosted by a nearly holocrystalline trachyte, is almost totally crystallized and shows textural characteristics comparable with those of the other enclaves, except for the amount of glass. The traces of glass in the 16B enclave have a felsic composition with SiO2 >70 wt %. (b) Several enclaves (e.g. 22x, 25x, 16bx) show radial intergrowths of Fe-augite and oligoclase plagioclase, which formed before anorthoclase crystallization. Such a radial texture of clinopyroxene and plagioclase in skeletal intergrowth (Fig. 3 g), although unusual, has been reported previously in medium-grained dolerite (MacKenzie et al., 1982). The presence of euhedral plagioclase at the glass contact seems in contrast with the xenomorphic intergrowth with the clinopyroxene. (c) Acicular apatite appears to intersect different crystals of plagioclase and the intracrystalline glass (Fig. 3e and f). This fragile texture (by SEM imaging we observed 1160 µm x 1 µm dimensions) is not compatible with progressive crystallization of magma in the residual spaces between earlier formed crystals in a partially molten/plastic enclave. (d) The glass in the enclaves is inhomogeneous, varying in colour from dark brown to clear over short distances (Figs 3 and 4) between the part in contact with the crystals and the central area, but, nevertheless, in some cases has a constant major element composition (Table 10). Alternatively a glass area that appears homogeneous in colour can show inhomogeneity in composition. The colour changes of interstitial glass have been used as a criterion to identify areas of melt-mixing within enclaves (Zinngrebe & Foley, 1995). (e) A crystal mush according to Marsh (1988) can form when just 50% of the melt has crystallized. The glass composition is, therefore, incompatible with the presence of a large amount of crystalline solid (also >90%) (Table 10) and with a residual felsic melt, which solidified as a glass instead of forming a granophyric texture.
(2) The origin of glass in the enclaves by host magma infiltration does not fit with the observed glass inhomogeneity (Table 10). For example, the Sr content of the glass can be very variable, even in adjacent point analyses, from >1000 ppm to <40 ppm (below detection limit). The high contents of Sr, not shown, are clearly due to feldspar melting in the enclave, whereas the very low Sr contents are due to host melt infiltration: the lack of mixing between two similar melts can be due to the greater viscosity of the host magma with abundant microlites and kinetic considerations. The host trachyte magma tends to fracture, permeate and dismember the peripheral part of the enclave (Fig. 2d).
(3) In situ partial melting appears to be most realistic explanation for the presence of interstitial glass. In medium-grained enclaves needles of apatite are enclosed in both oligoclase and anorthoclase: the melting of the anorthoclase leaves the residual apatite needles and partially preserves the plagioclase. Apatite is often the only crystal in the intracrystalline glass. The glass can occasionally show square-shaped K-feldspar microlites, which are portions of mechanically dismembered anorthoclase (Table 8). The high Sr contents of the glass are clearly due to feldspar melting (e.g. 1500–5800 ppm Sr in feldspar of 22x). The glass with the highest silica content (SiO2 >70%) occurs in the enclaves hosted within trachyte and, following Harris & Bell (1982), can be interpreted as the result of partial remelting of the majority of the low-temperature minerals, including quartz (no longer present). During this process the euhedral cores of more Ca-rich plagioclase crystals have been preserved. A least-squares mass-balance model, applied to the glass of sample 22, can be used to test this remelting hypothesis. The results are in excellent agreement with the hypothesis of partial melting of all the original low-temperature minerals in the enclave, particularly anorthoclase, with a contribution from late and interstitial minerals such as quartz and aegirine–augite, in essentially oxidizing conditions (Table 13). The calculated composition of the melted feldspar (Ab56An4Or40) is similar to the alkali-feldspar (Or38) found in syenite and in the recrystallized trachyte groundmass (Fig. 3 h). The Ab–Or–Qz haplogranitic system at 0·2 GPa shows that, after preliminary eutectic consumption of a small quantity of quartz, the melting process involves only the alkali feldspars. However, the high melting rate probably prevented the achievement of equilibrium conditions, leading to the formation of an inhomogeneous glass. In the Q–Ab–Or system (e.g. Bowen & Tuttle, 1950; Carmichael & MacKenzie, 1963) the glasses in the enclaves hosted by trachyte (excluding the glass in dolerite 21x) plot between the eutectic minimum (pantelleritic glass) and the thermal valley towards Or38 in the Or–Ab binary system (trachytic glass).
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, unreported data. H2O (wt %) content and temperature are reported as pairs of numbers. The large circle on the Ab–Or side of the diagram indicates the composition of the calculated felsic end-member with trachyte-number = 100% (
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Melting conditions of the enclaves
As noted above, the presence of glass in the enclaves is strictly related to the nature of the host magma: glass occurs only when this is trachytic. The Pantelleria trachytes have an estimated liquidus temperature of 950–1000°C (Mahood, 1984
; Lowenstern et al., 1993); similar temperatures were obtained by applying the liquid geothermometer of Sisson & Grove (1993b) and considering for the original melt a H2O content of 0·5–1·85 wt %. For the internal part of a solid enclave, under dry conditions and for a pressure range of 0·5–0·2 GPa, partial melting starts at 925–950°C. The Sisson & Grove (1993b) geothermometer, initially calibrated for liquids with SiO2 <55 wt % and not alkalic composition, is applicable in the range 1300–1100°C, under dry conditions, and 1100–750°C under wet conditions. The estimated temperatures of formation for the silica-saturated glasses (SiO2 64–66 wt % and AI 0·94) in the enclaves are almost 10–15°C lower than those of the host magma, 876°C for the high-silica glass (SiO2 >70 wt %), considering the maximum H2O content of the melts to be close to that of the host magma. The pressure of a column of melt with a density d = 2·5–2·2 g/cm3 in a magma chamber at 5 km depth (Mahood, 1984) falls within the range of 0·13–0·11 GPa; therefore, on the basis of the system Qz–Or–Ab–H2O, the initial melt of the enclaves at 850°C should have a minimum water content of 1·5 wt % (Johannes, 1985; Johannes & Holtz, 1990) and a water solubility limit of 4·2 wt % (Behrens & Jantos, 2001). The high-silica glass probably corresponds to incipient remelting, in conditions of partial equilibrium, whereas the other glasses result from the greater remelting of syenodioritic and syenogabbro enclaves in essentially H2O very undersaturated conditions (Fig. 14).
The residence time of the enclaves within the melt
According to Geschwind & Rutherford (1995), the presence in a melt of acicular and skeletal microlites or fibrous terminations of microphenocrysts is due to rapid decompression of the magma in the higher parts of the volcanic conduit, and to the high rate of crystallization (2–9 days). The presence of two different kinds of glass with similar composition, but containing rare and abundant microlites, respectively, in the peripheral portions of the enclaves is considered evidence of different residence times of the microlites within the melt, certainly longer in the host melt.
Feldspars from most of the enclaves in the trachytes of the ‘Cinque Denti’ caldera are zoned, but the anorthoclase crystals do not show exsolution. Therefore, the temperatures reached in the enclaves must have been high enough to account for the partial re-homogenization of pre-existing perthites. According to Brown & Parsons (1984), ‘only extremely rapidly chilled feldspars will remain homogeneous on cooling’ (at rates of a few degrees/hour). For zoned anorthoclase-bearing hornfels xenoliths in the basaltic andesite lavas of Stromboli (Aeolian Islands, Sicily), a cooling rate of a few days has recently been estimated (Renzulli et al., 2003).
As far as the zonation is concerned, the alkali feldspars have ovoid cores with sharp boundaries against the surrounding rims. Taking into account feldspar interdiffusion, which according to Christofersen et al. (1983) is not dependent on pressure and the H2O content of the melt, we have tested the residence time of the enclaves in the trachytic magma at 1000°C, on the basis of the relation D = 
2/(4t), where D is the interdiffusion rate (cm2/s) of a sphere of (cm) radius at t seconds (Johannes, 1985). Alkali feldspars have interdiffusion rates that vary with composition and crystallographic direction (Christofersen et al., 1983); for a value of D = 10–12 cm2/s the homogenization rate is approximately 1 µm/h, 10 µm in 3 days or 1 mm in 79 years. Plagioclase, according to the D values calculated by Johannes (1985) of 10–13 cm2/s at 1000°C, requires 10 times longer: 1 µm in 7 h, 10 µm in 29 days or 1 mm in 792 years (Fig. 15).
The interdiffusion rate within the (Or40) alkali feldspar at the pantellerite liquidus temperature of 850°C is approximately D = 10–13·5 cm2/s (Christofersen et al., 1983), corresponding to a homogenization rate of 1 µm/day or 10 µm in 91 days (Fig. 15). Mesoperthite exsolution and evidence of cataclastic deformation induced by thermal shock (Ferla et al., 1978) in the alkali feldspar in syenite enclaves within the host pantellerite are still preserved under these conditions.
On the basis of the above data, the residence time of the enclaves within the host magma is considered to be fairly short. Assuming thermal equilibrium, the lack of homogenization of the anorthoclase, on the basis of microscopical observations and electron microprobe analyses of sector zoning 10 µm thick, implies that the residence time of the enclave within the melt at 1000°C can be estimated at 3–5 days, if we consider the range of feldspar compositions and a linear interdiffusion rate based on extrapolation of previous data.
The ‘Daly Gap’ at Pantelleria
The ‘Daly Gap’ at Pantelleria between the mafic and felsic volcanic rocks has been explained by a number of models including crustal melting induced by basalt heating (Rittmann, 1967). Crustal melting is not confirmed by the relatively constant 87Sr/86Sr of both mafic and felsic volcanic rocks, which suggests that the felsic magmas could be derived from parental basalts by fractional crystallization (Barberi et al., 1969; Civetta et al., 1984, 1998). The latter hypothesis would be confirmed by the presence of enclaves of intermediate composition in the lavas (Villari, 1974; Civetta et al., 1984). This model has been successfully applied to explain the petrogenesis of bimodal alkaline volcanic suites in a number of continental rifts (Barberi et al., 1975; Peccerillo et al., 2003) and oceanic islands such as Mururoa (Caroff et al., 1993). In alternative hypotheses, pantellerites were thought to be derived from basalt unmixing (Lucido, 1982) or partial melting of a metasomatized mantle source (Bailey & Macdonald, 1987). Nevertheless, on the basis of isotopic studies, Mahood et al. (1990) have suggested for Pantelleria that trachypantellerite magmas were generated by partial fusion (5%, on the basis of mass-balance calculations; Ferla & Meli, in preparation) of crustal alkali-gabbros induced by intrusion of fresh basalt. The subsequent differentiation of the trachypantellerite magma resulted in the formation of anorthoclase-cumulate trachytes and residual pantellerite (Mahood & Baker, 1986; Mahood & Stimac, 1990; Mahood et al., 1990; Lowenstern & Mahood, 1991). According to Mahood & Baker (1986) the ‘benmoreitic’ enclaves in the Pantelleria volcanic rocks are feldspathic cumulates and are not representative of magma of intermediate composition in the ‘Daly Gap’.
Magma mixing and mingling can be significant processes in the generation of rocks of intermediate composition in the ‘Daly Gap’ of alkaline magmatic suites. However, although the chemical composition of the enclaves fills the ‘Daly Gap’, this does not indicate that there was ever a continuous liquid line of descent from mafic to felsic magma compositions.
The intermediate composition enclaves exhibit microscopic evidence for several generations of minerals that can either be rimmed or occur as single crystals: these include olivine, clinopyroxene, plagioclase, alkali feldspar and Fe–Ti oxides. The presence of forsteritic olivines in the enclaves, which are clearly xenocrystic, is consistent with the presence of mafic enclaves within a more felsic liquid. Normally zoned feldspars coexist with crystals with corroded cores of ternary feldspar, external euhedral part of the cores composed of plagioclase and outer zones of alkali-feldspar. According to Feeley & Dungan (1996), these textures can be explained by the inclusion of Na-rich crystals within a hotter, Ca-rich, melt, followed by crystallization of plagioclase from the hybridized magma, and eventually by the intercumulus-like crystallization of anorthoclase and alkali-feldspar from the residual melt.
Consequently we consider that the mineralogy of the enclaves is consistent with commingling between mafic and felsic magmas (i.e. Ferla, 1997). The more mafic enclaves are clearly highly differentiated mafic melts (hawaiite) containing relict ‘basaltic’ minerals such as forsteritic olivine, low Ti-augite, labradoritic plagioclase and Fe–Ti oxides. Subsequently, the commingling between hawaiite and trachyte melt containing cumulus anorthoclase took place. We suggest that this process began with an initial mixing (= blending) between the prevailing mafic magma and the felsic melt [according to the model of Frost & Mahood (1987)] (e.g. sample 21x enclave) and continued with further crystallization of the mafic components (Dorais et al., 1990; Wiebe et al., 1997) and incipient melting of the anorthoclase, as observed in the doleritic enclaves of hawaiitic composition.
The enclaves of intermediate composition, containing xenocrysts of the earlier stages of hawaiite and trachyte crystallization, provide the evidence for the commingling process, and undercooling inducing the crystallization, from an already hybrid melt, of intermediate composition acicular and skeletal olivine, Al- and Ti-rich rimmed pyroxene, radial intergrowths of pyroxene and oligoclase and acicular apatite. Late-stage crystallization of ternary feldspar, alkali-feldspar, quartz, aegirine–augite and, occasionally, Na-amphibole and biotite from the residual melt occurred in intercumulus-like conditions. The undercooling appears to be induced by the lower temperature of the original trachytic magma.
The mafic trachyte composition of the volcanic rocks hosting these enclaves falls within the linear array between the differentiated basalt and the trachyte in Harker variation diagrams. According to field and petrographic data, the host-rocks immediately surrounding the intermediate composition enclaves appear to be partially contaminated, enclosing solid and partially melted dismembered enclaves of small size, in agreement with the scheme proposed by Bacon & Metz (1984), Bacon (1986) and Feeley & Dungan (1996). These contaminated trachytes are the result of the mixing between fragmented intermediate enclave materias and a low-silica trachyte (Figs 2d and 6).
According to Mahood et al. (1990), Pantelleria provides proof of the existence of a stratified magma chamber with the pantellerite magma localized in the higher part, the anorthoclase-rich trachytic magma in the middle part and the basaltic magma localized below the trachyte in the lower part (Civetta et al., 1984, 1988, 1998; Mahood & Hildreth, 1986). On the basis of the results presented here, an alternative hypothesis to the model of continuous magmatic differentiation from basalt through trachyte to pantellerite can be suggested for the Pantelleria magmatic suite. Following Bacon & Metz (1984) and Bacon (1986), we propose that the enclaves with intermediate compositions demonstrate the interaction between the lower, not basaltic but differentiated mafic magma (hawaiite) and the higher trachyte containing anorthoclase, with the formation of a hybridized layer.
Some of the enclaves result from the intrusion of a fresh batch of basaltic magma into the more felsic part of a stratified magma chamber. The medium-grained enclaves have a more complex origin. The medium grain size of the enclaves, together with the presence of interstitial melt induced by heating from the host trachytic magma, most probably reflects fragmentation of the wall-rocks to the magma-chamber during the eruption of the ‘Green Tuff’ and accompanying caldera collapse. According to this hypothesis these enclaves would correspond to earlier crystallization products of the magma chamber system. During the ‘Green Tuff’ eruption, Pantelleria was covered by a large effusion of pantelleritic ignimbrite, followed by subordinate trachyte lavas (Civetta et al., 1998; Mahood & Stimac, 1988). The emptying of the magmatic chamber, the collapse of the upper part of volcanic edifice, and the formation of Cinque Denti caldera were the result of this eruption. Volcanic activity continued, according to Mahood & Hildreth (1986), with the intra-caldera effusion of the trachytic lavas hosting the enclaves, which form the basis of this study.
Partial melting of some of the enclaves in contact with host magma at a temperature of 1000–950°C occurred under very low-pH2O conditions, at pressures corresponding to a depth of between 2 and 5 km. Their residence time within the rising magma must have been less than 5 days, on the basis of the preserved zoning of the minerals. This is the calculated time for a magma rising from 2 km depth from the upper part of the chamber at a rate of approximately 12 m3/s through a conduit radius of 25 m (Wallmann et al., 1988).
Wall-rock enclaves, similar to those analysed in this study, should also occur within the intra-caldera ignimbritic pantellerites of more recent magmatic cycles. This appears to be the case for the 31x syenodioritic enclave and the ‘benmoreitic’ sample reported by Avanzinelli et al. (2004), which were both collected at the same locality (C.da Mida), within VI cycle pantellerite host-rocks.
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APPENDIX: SAMPLE LOCATIONS |
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TOPABSTRACTINTRODUCTIONPETROGRAPHIC DESCRIPTIONSDISCUSSION OF MAGMA EVOLUTIONAPPENDIX: SAMPLE LOCATIONSREFERENCES |
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ACKNOWLEDGEMENTS |
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The authors wish to thank C. Garbarino for his kind assistance with microprobe analysis at CNR laboratory of Cagliari (Italy), and S. Rotolo and P. Censi for helpful discussions. We are very grateful to C. W. Devey, F. Innocenti, G. A. Mahood, D. F. Parker, A. Provost, M. Stein, M. Wilson and two anonymous reviewers for constructive critical review and useful suggestions that greatly improved the manuscript. Financial support was provided by Italian ‘Ministero Istruzione Università Ricerca’ (MIUR, 60%).
*Corresponding author. Telephone: +39 091 6161516. Fax: +39 091 6168376. E-mail: pferla{at}unipa.it, applina2{at}katamail.com
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TOPABSTRACTINTRODUCTIONPETROGRAPHIC DESCRIPTIONSDISCUSSION OF MAGMA EVOLUTIONAPPENDIX: SAMPLE LOCATIONSREFERENCES |
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