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Journal of Petrology Volume 42 Number 10 Pages 1813-1843 2001
© Oxford University Press 2001
Almandine Garnet in Calc-alkaline Volcanic Rocks of the Northern Pannonian Basin (Eastern–Central Europe): Geochemistry, Petrogenesis and Geodynamic Implications
1DEPARTMENT OF PETROLOGY AND GEOCHEMISTRY, EÖTVÖS UNIVERSITY, H-1088 BUDAPEST, MÚZEUM KRT. 4/A, HUNGARY
2SCHOOL OF EARTH SCIENCES, BIRKBECK COLLEGE, MALET STREET, LONDON WC1E 7HX, UK
3DEPARTMENT OF GEOLOGY, ROYAL HOLLOWAY UNIVERSITY OF LONDON, EGHAM TW20 0EX, UK
4PETROLOGY GROUP, FACULTY OF EARTH SCIENCES, UNIVERSITY OF UTRECHT, BUDAPESTLAAN 4, 3584 CD UTRECHT, THE NETHERLANDS
Received August 18, 2000; Revised typescript accepted March 9, 2001
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONGEOLOGICAL BACKGROUNDANALYTICAL TECHNIQUESALMANDINE GARNETSGEOCHEMISTRY OF GARNET-BEARING...DISCUSSIONREFERENCES |
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Almandine garnet-bearing andesites and dacites occur frequently in the Neogene calc-alkaline volcanic series of the northern Pannonian Basin (Hungary and Slovakia). They were erupted during the early stage of volcanism and occur along major tectonic lineaments. On the basis of petrographic and geochemical characteristics, garnets from these rock types are classified into (1) primary phases, (2) composite minerals containing xenocrystic cores and magmatic overgrowths and (3) garnets derived from metamorphic crustal xenoliths. Coexisting phenocrysts of primary garnets include Ca-rich plagioclase, hornblende (magnesiohastingsite to tschermakite) and/or biotite. The primary garnets have high CaO (>4 wt %) and low MnO contents (<3 wt %). They have strongly light rare earth element depleted patterns and are enriched in heavy rare earth elements. Negative Eu anomalies occur only in garnets in the more silicic host rocks.
18O values for primary garnets are 6·1–7·3
, whereas composite garnets have elevated 18O values (>8). Chemical compositions of the primary garnets and coexisting minerals suggest that they crystallized at high pressures (7–12 kbar) and temperatures (800–940°C) from mantle-derived magmas. Sr–Nd isotopic compositions of their host rocks and O isotopic values of the garnets are consistent with two-component mixing between mantle-derived magma and lower-crustal metasedimentary material. The garnet-bearing silicic magmas were erupted during extension of the Pannonian Basin and the tensional stress field may have enhanced their fast ascent from lower-crustal depths, allowing preservation of early-formed almandine phenocrysts. KEY WORDS: almandine; garnet; calc-alkaline volcanism; geochemistry; Pannonian Basin
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDANALYTICAL TECHNIQUESALMANDINE GARNETSGEOCHEMISTRY OF GARNET-BEARING...DISCUSSIONREFERENCES |
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Almandine garnet is a relatively rare mineral phase in calc-alkaline volcanic rocks worldwide (Table 1). Its rarity may be due to the restricted conditions under which such garnets can form (hydrous mantle source, high-pressure crystallization from hydrous Al-rich magma) and to the particular geodynamic setting (tensional stress field), which enhances the rapid ascent of garnet-bearing melts. Green & Ringwood (1968)
demonstrated that almandine-rich garnet could be a liquidus or near-liquidus mineral in silicic magmas at high pressure (9–18 kbar). Further experimental studies (e.g. Hensen & Green, 1973; Green, 1976, 1977, 1982, 1992; Clemens & Wall, 1981; Conrad et al., 1988) showed that Ca-rich (CaO >4 wt %) and Mn-poor (MnO <4 wt %) almandine could crystallize from hydrous andesitic magmas at relatively high pressure (>7 kbar) and at temperatures of 900–950°C. Green (1977) suggested that the grossular and spessartine contents of almandine are sensitive to pressure–temperature conditions. The grossular content of garnet increases with increasing pressure, whereas a higher Mn content stabilizes garnet at shallower depths, so that Mn-rich (MnO >4 wt %) almandines can crystallize from silicic liquids at 
5 kbar. Magma composition exerts a strong influence on garnet chemistry. Garnets hosted by peraluminous (S-type) volcanic rocks are CaO poor (CaO <4 wt %), whereas garnets crystallized from metaluminous (I- or M-type) magmas have higher CaO content (CaO >4 wt %). In summary, the chemistry of almandine-rich garnet is sensitive to the conditions of formation, indicating the petrogenetic importance of this mineral in volcanic rocks.
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Ca-bearing almandine garnet is unstable at low pressure (e.g. Green, 1982
, 1992). At shallow depths resorption of garnet begins and garnet may disappear during a long residence time in a shallow magma chamber. Therefore, the presence of primary Ca-bearing almandine in volcanic rocks suggests relatively rapid ascent of the host magma (e.g. Fitton, 1972; Gilbert & Rogers, 1989), which could be enhanced by a tensional stress field (regional extension).
Almandine occurs in Neogene calc-alkaline volcanic rocks of the Carpathian–Pannonian region of Eastern–Central Europe (Fig. 1a) and is relatively common in the Northern Pannonian Basin (Hungary and Slovakia; e.g. Brousse et al., 1972; Embey-Isztin et al., 1985; Lantai, 1991; Szabó et al., 1999). The geodynamic setting and the relationship between volcanism and tectonic events in this region have long been the subject of debate (e.g. Lexa & Kone
ný, 1974; Balla, 1981; Szabó et al., 1992; Lexa, 1999; Harangi & Downes, 2000). The rarity of garnet-bearing volcanic rocks in general and the experimental results pointing to their high-pressure origin suggest that their formation and eruption require particular geodynamic conditions. Therefore, garnet-bearing volcanic rocks could carry useful information concerning the geodynamics of magma genesis in this area. In addition, a detailed study of the garnets and their host volcanic rocks in the Carpathian–Pannonian region will contribute to our knowledge of the genesis of garnet-bearing magmas worldwide.
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In this paper we present a petrologic and geochemical study of garnets and their host rocks from the Northern Pannonian Basin (NPB). We provide major and trace element data for the garnets and the first oxygen isotope data for igneous almandines. Interpretation of the major and trace element composition and Sr and Nd isotopic ratios of the host rocks, combined with the oxygen isotope composition of the garnets, suggests a petrogenesis involving mantle-derived host magmas and lower-crustal contamination processes. We conclude that the eruption of garnet-bearing magmas within the Pannonian Basin was related to an extensional tectonic phase.
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GEOLOGICAL BACKGROUND |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDANALYTICAL TECHNIQUESALMANDINE GARNETSGEOCHEMISTRY OF GARNET-BEARING...DISCUSSIONREFERENCES |
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Calc-alkaline volcanism along the Carpathian volcanic arc took place from the Miocene to the Quaternary. Although a direct relationship has been suggested between volcanism and subduction in the East Carpathian arc (e.g. Bleahu et al., 1973
; Boccaletti et al., 1973; Mason et al., 1998), genesis of the calc-alkaline volcanic rocks of the West Carpathians is more controversial. A subduction-related genesis was proposed by Balla (1981), Szabó et al. (1992) and Downes et al. (1995), whereas Lexa & Konený (1974), Lexa et al. (1993), Harangi et al. (1999) and Lexa (1999) suggested that melt generation was due to extension of the Pannonian Basin.
The calc-alkaline volcanic rocks of the Carpathian arc are dominantly andesites and dacites with subordinate rhyolites and basaltic andesites (e.g. Szabó et al., 1992; Kaliiak & 
ec, 1995; Konený et al., 1995; Mason et al., 1996). Garnet-bearing volcanic rocks occur throughout the arc, but are most abundant in the west (Fig. 1a). This segment overlaps the northern part of the Pannonian Basin and comprises the Central Slovakian Volcanic Field (CSVF), the volcanic complexes of the Börzsöny and Visegrád Mts and the Cserhát and Mátra volcanic areas (Fig. 1b). Volcanism in the NPB started with eruption of medium-K to high-K garnet-bearing magmas (16·5–16 Ma; Konený et al., 1995; Harangi, 1999), followed by the eruption of garnet-free, dominantly andesitic volcanic products (Konený et al., 1995; Harangi & Downes, 2000). The garnet-bearing volcanic rocks are situated along major tectonic lines, such as the Pilis line in the southwestern Visegrád Mts, the Diósjen
line across the Börzsöny Mts, the Etes tectonic zone in the Karancs and the Central Slovakian Fault across the CSVF (Fig. 1b). Calc-alkaline volcanism terminated at 
9 Ma (Pécskay et al., 1995) and was followed by alkaline basaltic magmatism within the CSVF and in the Nógrád–Gömör Volcanic Field (Fig. 1b). The crust and lithosphere beneath the NPB thin gradually towards the south. The relatively thin crust (28–32 km) and lithosphere (70–90 km; Horváth, 1993) may imply that evolution of this area was influenced by the Mid-Miocene syn-rift extension of the Pannonian Basin.
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ANALYTICAL TECHNIQUES |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDANALYTICAL TECHNIQUESALMANDINE GARNETSGEOCHEMISTRY OF GARNET-BEARING...DISCUSSIONREFERENCES |
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Electron microprobe data were obtained using a JEOL-733 Superprobe equipped with an Oxford Instrument ISIS energy dispersive system (Birkbeck College, University of London) using 15 kV accelerating voltage for 100 s. Some garnet analyses were also carried out using a JEOL Superprobe JXA-8600 wavelength dispersive system at the University of Florence; compositions obtained by both methods were the same within analytical error.
Selected garnets were analysed for their O isotope composition at Royal Holloway University of London (RHUL) using a laser fluorination system attached to a VG Isotech PRISM gas-source mass spectrometer (Mattey & Macpherson, 1993). Samples (1·4–1·7 mg; 7–10 grains) of inclusion-free garnet were analysed, usually in duplicate. The samples were heated with a Nd–YAG laser in the presence of ClF3. Yields were >92% (mostly >95%) for all the samples. San Carlos (SC) olivine was used as a secondary standard. Data, recorded as 18O values, are given with reference to Standard Mean Ocean Water (SMOW). Details of accuracy and precision have been given by Mattey & Macpherson (1993).
In situ trace element compositions of the same garnets were obtained by laser ablation inductively coupled plasma mass spectrometry (ICP-MS) at the University of Utrecht using a 193 nm ArF excimer laser ablation system (Geolas, Microlas; Günther et al., 1997) in combination with a quadrupole ICP-MS instrument (Platform ICP, Micromass). The method has been described in more detail by Mason et al. (1999). Ablation was performed at a fixed point on the sample with a fluence of 20 J/cm2, a laser shot repetition rate of 10 Hz and an ablation crater diameter of 120 µm, resulting in an ablated volume of 1–2 µg. Calibration was performed against NIST SRM 612 glass using the concentration data reported by Pearce et al. (1997). Calcium, previously determined by electron microprobe analysis at the surface of the volume removed by laser ablation, was used as an internal standard. The first 5–10 µm of material removed from the sample was not used in concentration calculations to remove the effect of surface contamination. Rare earth element results for ‘in-house’ garnet standards were in good agreement (better than 10% for the most homogeneous samples) with secondary ionization mass spectrometry data. Data for spots in which an inclusion had clearly been analysed were discarded.
Major and trace element compositions of the whole rocks were determined on a Philips PW1480 X-ray fluorescence spectrometer at RHUL using fused glass discs and pressed powder pellets. Details of precision have been given by Baker et al. (1997). Rare earth elements (REE) were analysed by ICP-MS at RHUL using the method described by Walsh et al. (1981). REE were separated using ion exchange columns, after fusion and dissolution of samples.
Sr and Nd isotopic compositions of the whole rocks were determined on a VG354 multicollector thermal ionization mass spectrometer at RHUL. Sample powders for Sr isotope analysis were leached in hot 6 M HCl for 1 h to remove any effects of postmagmatic groundmass alteration. Sr and Nd isotope ratios were determined in multidynamic mode as described by Thirlwall (1991). 87Sr/86Sr was normalized to 86Sr/88Sr = 0·1194 and 143Nd/144Nd to 146Nd/144Nd = 0·7219. An in-house laboratory Nd standard (Aldrich) yielded 143Nd/144Nd = 0·511412 ± 5 (2 SD on eight analyses; equivalent to 0·511847 for the international La Jolla standard) and the SRM987 gave 87Sr/86Sr = 0·710241 ± 16 (2 SD on 16 analyses).
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ALMANDINE GARNETS |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDANALYTICAL TECHNIQUESALMANDINE GARNETSGEOCHEMISTRY OF GARNET-BEARING...DISCUSSIONREFERENCES |
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Petrography and classification
Almandine garnet occurs in a wide variety of volcanic rocks (andesites to rhyodacites) of the NPB and also in rare crustal xenoliths within the volcanic rocks. Garnet does not exceed 1–2 vol. % in abundance and is homogeneously distributed in each rock type. The garnets are dark red–violet and sometimes are surrounded by a thin light-coloured corona. They are usually 1–8 mm in diameter, but some may be as large as 2 cm. In general, garnets are larger in andesitic host rocks than in the more silicic ones.
The texture of the host rocks is variable, ranging from sparsely porphyritic trachytic to strongly porphyritic and medium grained. Plagioclase is the most common phenocryst coexisting with garnet. Hornblende and biotite occur also frequently and are together in the most silicic andesites and less silicic dacites. Biotite is the only mafic coexisting mineral in the rhyodacites but is absent from the andesites. Orthopyroxene occurs sporadically as a microphenocryst or groundmass phase, whereas clinopyroxene has never been observed. Mineral phases characteristic of many garnet-bearing S-type volcanic rocks worldwide, such as quartz, cordierite and K-feldspar, are absent from the NPB garnet-bearing volcanic rocks. On the basis of the petrographic characteristics of the garnets and their host rocks, the NPB almandines can be classified into five groups. Type 1A, Type 1B and Type 2 garnets are homogeneous or slightly zoned crystals with an inclusion assemblage similar to the mineral phases occurring in the host rocks. They appear to be primary igneous phases. Type 3 garnets are composite crystals having a light-coloured xenocrystic core and magmatic overgrowth. Type 4 garnets occur in crustal xenoliths. Details of petrographic and geochemical characters of the garnet types are given in Table 2.
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Type 1A almandines (Fig. 2a and b) are 1–4 mm in size and occur in glassy rhyodacites (SiO2 69–71 wt %) of the Visegrád Mts. They are euhedral to subhedral, but occasionally slightly rounded. Flow lines can be often observed around them (Fig. 2a). They are generally inclusion poor, the most common inclusions being plagioclase and needles of apatite. Intergrowth with plagioclase and/or biotite can be observed. Type 1B almandines (Fig. 2c and d) are the common garnets in silicic andesites and dacites (SiO2 62–67 wt %) of the Börzsöny Mts. They are usually 2–5 mm in diameter, typically subhedral and often rounded (Fig. 2c). Intergrowths with plagioclase or biotite (Fig. 2c and d) and rarely with hornblende are characteristic. The almandines contain variable amounts of inclusions, mostly plagioclase, ilmenite and apatite. Type 2 garnets (Fig. 2e and f) occur in andesites (SiO2 58–62 wt %) of Karancs and Breziny. They are usually 3–10 mm in diameter, but can reach 1·5 cm. They are always rounded and have resorbed, irregular margins. Inclusions are abundant and consist of mostly plagioclase, often in a concentric arrangement (Fig. 2f).
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Composite garnets (Type 3; Fig. 2g and h) have been found in the Börzsöny Mts and in the Karancs (Fig. 1b). They are characterized by an irregular light-coloured core surrounded by a typical Type 1B or Type 2 outer garnet zone. Resorption between the light-coloured core garnet and the dark mantle garnet can often be observed. The core of some composite garnets has a distinct inclusion assemblage and/or contains xenoliths. Figure 2g shows an irregular-shaped crustal xenolith in the light-coloured core of a rounded garnet. This rock fragment consists of plagioclase, corundum, pleonaste (hercynite) and biotite. The garnet core includes further tiny fragments of this xenolith. A Type 3 garnet core from
iatoro
(Fig. 1b) contains inclusions of biotite, ilmenite and orthopyroxene characterized by relatively high MnO content.
Crustal xenoliths are common within the garnet-bearing volcanic rocks from iatoro and Breziny, but do not contain almandine. Almandine-bearing xenoliths occur rarely in the garnet-free volcanic rocks of the Börzsöny and Visegrád Mts. One of these samples consists of rounded garnets (Type 4; Fig. 2i) coexisting with plagioclase, biotite, sillimanite, pleonaste (hercynite) and orthopyroxene. The core of these garnets contains abundant quartz inclusions, but the outer zones are usually inclusion free.
Geochemistry of garnets
Major elements
Representative major element analyses of garnets from the NPB volcanic rocks are shown in Table 3. Comparisons of estimated Fe2O3 contents with values measured by Mössbauer spectrometry have revealed discrepancies particularly at low Fe2O3 content (e.g. Luth et al., 1990). Measured Fe2O3 content of eight garnet samples by Mössbauer spectrometry (Kósa, 1998) was between 1·25 and 1·45 wt %, whereas the estimated Fe2O3 in the same samples was significantly lower (0·55–1·16 wt %). In this study, we used a stoichiometric method (Muhling & Griffin, 1991) for ferric ion estimation. End-member molecules were calculated based on the method of Muhling & Griffin (1991). All the published garnet analyses used for comparison in this study were recalculated using the same scheme.
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The petrographically distinguished garnet types have distinct chemical compositions illustrated in the Ca–Mg–Fe diagram (Fig. 3) and in Table 2. All of the NPB garnets have 58–73 mol % almandine component. Type 1A garnets show remarkably homogeneous chemistry, whereas Type 1B and Type 2 garnets are more variable. Rhyodacite-hosted Type 1A garnets have the highest almandine component, whereas Type 1B garnets from dacites and Type 2 garnets from andesites have lower almandine contents (Table 2). Among the minor components, pyrope is in the range of 9–14 mol % in Type 1A garnets, 13–23 mol % in Type 1B garnets and 12–21 mol % in Type 2 garnets. In contrast to the rough similarity between Type 1B and Type 2 garnets in terms of Mg–Fe abundance, there is a significant difference in the grossular and spessartine components. Overall, almandines from the NPB volcanic rocks have relatively high grossular (Gro >6 mol %) and low spessartine (Spe <10 mol %) contents. Type 1B garnets are the richest in Ca (Gro 10–18 mol %) and have generally low spessartine component (Spe <4 mol %). Andesite-hosted Type 2 garnets have typically lower grossular (Gro 6–13 mol %) and higher spessartine content (Spe 4–10 mol %). Type 1A garnets from the most silicic host rocks have similar grossular (Gro 8–12 mol %), but lower spessartine (Spe 3·5–4·5 mol %) components.
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Single crystals are homogeneous or show slight zoning, whereas composite garnets (Type 3) have complex zoning patterns (Table 2). In Type 1A garnets a slight decrease of pyrope and slight increase of grossular from core to rim can be observed. Compositional trends in Type 1B garnets indicate increasing pyrope and increasing grossular from core to rim, whereas Type 2 garnets show mainly increasing pyrope and decreasing grossular and spessartine from core to rim.
Composite (Type 3) garnets have compositionally distinct cores with lower Ca contents (CaO <4 wt %, Gro <6·5 mol %; Fig. 3b) and higher Fe contents than the surrounding overgrowth. The MnO content of the cores is variable (0·5–6·5 wt %), but it is always higher than in the outer zone. The overgrowths show typical Type 1B or Type 2 compositions.
Type 4 garnet from the metamorphic xenolith is spessartine–almandine (60–64 mol % almandine, 18–25 mol % spessartine). Some of these garnets have a thin margin, in which Mn and Fe concentrations drastically decrease, whereas Mg and Ca increase.
Trace elements
Trace element compositions of selected garnets are presented in Table 4 and Fig. 4. They were determined by laser ablation ICP-MS analysing 2–4 points for each sample. The advantage of this method over the bulk garnet instrumental neutron activation analysis (INAA; e.g. Irving & Frey, 1978; Thirlwall & Fitton, 1983; Gilbert & Rogers, 1989) is that the influence of inclusions can be avoided, therefore it gives a better estimate of the trace element composition of garnets. However, we deliberately analysed a few garnet areas that contained inclusions such as plagioclase, apatite and zircon. These compositions enabled us to monitor and avoid the possible effect of these minerals on the garnet composition.
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All the NPB garnets show strong light REE (LREE) depletion (Ce/YbN = 0·001–0·005) and have typical heavy REE (HREE) enriched patterns with variable HREE concentrations (YbN 59–850; Fig. 4). The degree of LREE depletion is more pronounced in these garnets than in previous INAA-derived results (Irving & Frey, 1978; Thirlwall & Fitton, 1983; Gilbert & Rogers, 1989). The garnet types previously distinguished by their petrographic characteristics have different trace element compositions and different REE profiles. The total REE content and the Y, Zr, Hf, Th and U concentrations increase from Type 2 garnet through Type 1B to Type 1A garnets. Furthermore, the negative Eu anomaly tends to be stronger towards the Type 1A garnets (Table 2, Fig. 4). The core of the composite garnet (Type 3) from Verce (Börzsöny Mts) has lower Y and higher Zr, Hf concentrations than the primary garnets (Table 4) and shows a different REE pattern.
Oxygen isotopes
Although 18O values have been reported for metamorphic almandines (e.g. Taylor & Coleman, 1968; Kohn et al., 1997), our results are the first published data for igneous almandines since the single analysis of Mason et al. (1996). A large amount of laser fluorination oxygen isotope data for mineral phases from Neogene volcanic rocks of the Carpathian–Pannonian region has already been published (Downes et al., 1995; Mason et al., 1996; Dobosi et al., 1998). The samples were analysed in the same laboratory (RHUL) using the same technique (Mattey & Macpherson, 1993); therefore these data provide a comparative dataset.
Values of 18O for the NPB garnets show a wide range from 6·1 to 10·5 (Table 5). Type 2 garnets have the lowest 18O values (6·1 and 6·6), whereas values for Type 1A and Type 1B garnets are slightly higher but show a narrow range (6·9–7·3). Two composite garnets have elevated 18O values (8·0 and 10·5, respectively). Garnet in an andesite from the East Carpathians (Mason et al., 1996) has a slightly higher 18O value (7·3) than most of the Type 1A and 1B garnets. The oxygen isotope data for igneous garnets from the NPB do not correlate with the major element compositions of garnets, but correlate well with the host rock 87Sr/86Sr (Fig. 5) and 143Nd/144Nd values. 18O values of composite garnets are between the typical values of igneous almandines published in this study and metamorphic ones (11·4–13·4; Kohn et al., 1997), consistent with the assumption that both the xenocrystic core and igneous overgrowth of the composite garnet were analysed. Whole-rock data for garnet-bearing volcanic rocks from SE Spain and the Lesser Antilles (Table 1) show strongly elevated 18O values (13·1–15·6; Munksgaard, 1984; Davidson & Harmon, 1989). As whole-rock oxygen isotope data are usually higher than the included minerals by at least 1, garnets from these rocks could have higher 18O values than the NPB garnets. Garnets from the English Lake District (Table 1) have 18O values of 8·4–8·6 (M. F. Thirlwall, unpublished data, 1999), higher than those for the igneous garnets from the NPB.
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The oxygen isotope data of the NPB garnets are higher than the typical 18O values inferred for the asthenospheric mantle beneath the Pannonian Basin (5·0–5·4; Dobosi et al., 1998), but 18O values of Type 2 garnets are close to the values of the lithospheric mantle as shown by separated minerals of ultramafic xenoliths from the East Carpathians (18O = 5·4–5·9; Mason, 1995; Fig. 5).
Mafic minerals from the host volcanic rocks show roughly similar 18O values to the garnets, although some biotites have higher 18O values. Comparing the oxygen isotope data for the garnets with data for the mafic minerals (pyroxenes, hornblende) of the garnet-free volcanic rocks from the same area (Downes et al., 1995), only some amphiboles show lower 18O values (Fig. 5). In general, almandine garnets from the NPB have the lowest 18O values at a given host rock 87Sr/86Sr value among the calc-alkaline volcanic suite of the Carpathian–Pannonian region.
Inclusions and coexisting minerals
Hornblende is a common mafic phenocryst phase in the andesites and dacites, being intergrown with garnet in some samples. It shows variable resorption and formation of opacite rims. Hornblende compositions (Table 6) are tschermakite, magnesiohastingsite and pargasite based on the IMA nomenclature (Leake et al., 1997). Amphiboles intergrown with garnets have similar compositions to the phenocrysts. In general, there are positive correlations between Aliv and Ti, and Aliv and Na + K (Fig. 6), corresponding to crystallization under falling temperature and pressure (Gill, 1981; Green, 1992). Most amphiboles in the garnet-bearing volcanic rocks show higher Aliv, Ti and Na + K values than amphiboles of the garnet-free volcanic rocks and approach the values of the magnesio-hastingsites occurring in mafic cognate xenoliths from the NPB (Sz. Harangi, unpublished data, 1998).
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Biotite occurs in variable amounts in the garnet-bearing volcanic rocks of the NPB and is less common in the garnet-free varieties. The intergrowth and coexistence of biotite and garnet suggest contemporaneous crystallization. Biotite has distinct chemical compositions (Table 7) in the different rock types, illustrating compositional dependence on the nature of host magma (Abdel-Rahman, 1994). They are less magnesian (mg-number 0·29–0·40) and less potassic (K2O 8·1–8·7 wt %) in rhyodacites compared with micas in dacites and andesites (mg-number 0·45–0·53, K2O 8·5–9·4 wt %). The garnet–biotite Mg–Fe distribution coefficient, i.e. (Mg/Fe)gt/(Mg/Fe)bi, varies between 0·24 and 0·35. These values are higher than the Mg–Fe distribution coefficient typical of metamorphic garnet–biotite pairs (0·09–0·20; Lyons & Morse, 1970) but match that of the garnet–biotite pairs from igneous rocks (0·24–0·32; Lyons & Morse, 1970). The chemical composition of biotite in the garnet-bearing crustal xenolith is very different, having higher TiO2 (5·3 wt %) and MnO (0·74 wt %) contents.
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Plagioclase (Table 8) is the most common mineral in the garnet-bearing volcanic rocks of the NPB, occurring as phenocrysts, intergrowths with garnet (Fig. 2c), inclusions in mafic minerals and reaction coronas around garnet (Fig. 2e and f). The phenocrysts show normal, reverse and oscillatory zoning (An41–88). Plagioclase inclusions in garnet show the same chemical variation as the phenocrysts in each sample (An55–91; Fig. 7). Most of them are more calcic than the plagioclase inclusions in biotite and amphibole phenocrysts (An53–78).
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Ilmenite is a common minor inclusion in Type 1B and Type 2 garnets and occurs also as rare microphenocrysts. No ilmenite has been found in Type 1A garnets. Its composition (Table 9) differs in the two garnet types, with more TiO2 and less MnO in Type 1B garnets. Ilmenite in a xenocrystic core of a Type 3 composite garnet from the Karancs Mts contains significantly higher MnO (MnO 1·9–2·1 wt %).
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Orthopyroxene microphenocrysts are magnesian ferrosilite (En44–51Wo1·4Fs47–54). Orthopyroxene inclusions in garnet cores of the Karancs andesite contain relatively high MnO (MnO 2·3–5·9 wt %) coinciding with the higher MnO content of the garnet core (MnO >3·7 wt %) and the coexisting inclusion phases.
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GEOCHEMISTRY OF GARNET-BEARING VOLCANIC ROCKS |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDANALYTICAL TECHNIQUESALMANDINE GARNETSGEOCHEMISTRY OF GARNET-BEARING...DISCUSSIONREFERENCES |
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The garnet-bearing volcanic rocks in the NPB range in composition from andesites to rhyodacites (SiO2 58·2–70·1 wt %; Table 10). They show a continuous trend from medium- to high-K calc-alkaline volcanic rocks (Fig. 8) and are usually more silicic than the garnet-free volcanic rocks in the same region. Garnet-bearing rhyodacites are slightly peraluminous, having normative corundum in the CIPW composition (c = 1·5–2·1) and A/CNK ratio [Al2O3/(CaO + Na2O + K2O) ratio in molar proportions] > 1 (A/CNK = 1·07–1·12). Dacites containing Type 1B garnets are both peraluminous and metaluminous (c = 0–1·9, di = 0–1·7; A/CNK = 0·95–1·09), whereas Type 2 garnet-bearing andesites are diopside-normative (di = 0·6–1·6; A/CNK = 0·94–0·96).
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Trace element compositions of the garnet-bearing volcanic rocks (Table 10) are similar to those of the garnet-free varieties. They show relative enrichment of large ion lithophile elements (LILE), negative Nb anomalies and positive Pb anomalies in a mid-ocean ridge basalt (MORB)-normalized plot (Fig. 9), typical of subduction-related magmas. The REE show similar patterns to the garnet-free rocks in the range of La–Eu; however, the garnet-bearing rocks are relatively depleted in the HREE (Fig. 10). This HREE depletion is most pronounced in Type 1A and most Type 1B garnet-bearing dacites and rhyodacites. Similar strong HREE depletion was recognized in the garnet-bearing rhyolites from the Pyrenees (Gilbert & Rogers, 1989), whereas the English Lake District andesites and dacites (Thirlwall & Fitton, 1983) and several other garnet-bearing volcanic rocks worldwide (Irving & Frey, 1978) do not show this feature. The NPB andesites containing Type 2 garnets show no HREE depletion (Fig. 10). In general, the garnet-bearing volcanic rocks of the NPB do not show a significant negative Eu anomaly, in contrast to garnet-bearing dacites and rhyolites worldwide (e.g. Pyrenees, English Lake District, Japan, Victoria; Eu/Eu* = 0·32–0·74). The Eu/Eu* ratio varies between 0·76 and 0·97 in the NPB rocks, but most of them are in a narrow range from 0·87 to 0·97. These values are usually even higher than the Eu/Eu* ratios of the garnet-free volcanic rocks of the NPB at given LaN.
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Initial 87Sr/86Sr and 143Nd/144Nd ratios in the garnet-bearing volcanic rocks of the NPB vary widely (87Sr/86Sr = 0·7062–0·7100; 143Nd/144Nd = 0·51230–0·51246; Table 10) and reach the highest 87Sr/86Sr and lowest 143Nd/144Nd values among the calc-alkaline volcanic rocks in this area. The dacites and rhyodacites have fairly similar radiogenic isotope compositions (87Sr/86Sr = 0·7090–0·7100; 143Nd/144Nd = 0·51235–0·51246) and differ significantly from the garnet-bearing andesites (87Sr/86Sr = 0·7062–0·7069; 143Nd/144Nd = 0·51244–0·51246).
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DISCUSSION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL BACKGROUNDANALYTICAL TECHNIQUESALMANDINE GARNETSGEOCHEMISTRY OF GARNET-BEARING...DISCUSSIONREFERENCES |
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Origin of almandine garnets in the NPB volcanic rocks
Almandine garnets in volcanic rocks can be (1) primary magmatic phases or (2) xenocrysts derived either from disaggregated country rocks or represent a restite phase. The petrographic and geochemical characteristics of Type 1A, 1B and Type 2 almandines in the NPB volcanic rocks suggest that they are primary igneous phases cognate with their host rocks. Garnets are commonly intergrown with phenocrysts and contain an inclusion assemblage akin to the phenocrysts of the host rocks. Furthermore, plagioclase inclusions in the garnets and plagioclase phenocryst cores have identical compositions in each sample.
Garnets in the NPB andesites are usually surrounded by a reaction corona of plagioclase ± biotite or amphibole, whereas those in the more acidic rocks have sharp margins, euhedral or rounded shapes and lack reaction rims. The varying degree of resorption and the occurrence of reaction coronae may reflect differences in the speed of magma ascent. We infer that the rhyodacitic magma carrying the Type 1A garnets ascended and erupted rapidly. The rapid ascent prevented reaction and preserved the euhedral garnet shapes. The low amount of phenocrysts and the glassy groundmass of these rocks also support this idea. In contrast, the andesitic magmas could have had a slower ascent rate, so breakdown and resorption of early-formed Type 2 garnets took place at lower pressure.
The cores of composite garnets (Type 3) are in turn unambiguously xenocrysts, on the basis of their different chemical composition (low CaO and variable MnO contents) and their distinct inclusion assemblage (e.g. MnO-bearing biotite and orthopyroxene) or their enclosed metamorphic rock fragments (Fig. 2g). In the NPB, garnet occurs in different metamorphic rocks (amphibolite, gneiss, micaschist) within the crystalline basement, but these garnets have clearly different compositions compared with the cores of the composite garnets (Koroknai et al., 1999). Almandines in the gneisses are pyrope rich (Pyr 17–19 mol %), whereas they have a significant grossular component (Gro 18–20 mol %) in the amphibolites. In addition, K/Ar and Ar/Ar ages as well as zircon fission-track data indicate rapid extensional unroofing during the Late Cretaceous (Koroknai et al., 1999). Thus, these rocks were situated at shallow crustal levels at the time of the calc-alkaline volcanism (Middle Miocene). As the low-Ca Type 3 garnet cores are overgrown by primary garnet zones inferred to have crystallized at high pressure, these shallow crustal metamorphic rocks are unlikely to be a potential source for the xenocrystic garnet cores. Almandine garnets from the crustal xenolith in an andesite of the Visegrád Mts (Type 4 garnet) have a very different chemical composition (high MnO content) compared with the xenocrystic Type 3 garnet cores (Tables 2 and 3). Therefore, this metamorphic rock fragment also cannot represent the source rock of the low-Ca garnet cores. The crustal xenolith enclosed by the light-coloured core of a rounded garnet phenocryst in the Nógrád dacite (Fig. 2g), however, may provide some information about the source of the low-Ca garnets. This rock fragment consists of corundum, pleonaste and biotite and represents an Al-rich composition. Thus, this xenolith-bearing Ca-poor garnet may represent a restite formed by anatexis of a lower-crustal metapelite. The REE pattern of the light-coloured core of a Type 3 garnet has stronger negative Eu anomaly and higher Gd/DyN ratio than the primary garnet types (Fig. 4) and resembles that of garnets from granulite-grade metapelites (e.g. Bea et al., 1997; Bea & Montero, 1999). The irregular contacts and incomplete resorption between the garnet xenocrysts and the garnet overgrowth indicate reactions between the incorporated garnet fragments and the ascending melt.
Primary igneous garnets have great petrogenetic significance, because their composition depends on the magma type and the pressure and temperature under which they crystallize (Green, 1977, 1992). High-pressure almandines from mantle-derived (M-type) or I-type magmas (Table 1) are characterized by relatively high CaO (>5 wt %) and low MnO (<2 wt %) contents (Fig. 11a). With decreasing pressure, MnO content increases (>3 wt %), whereas CaO remains relatively high (>4 wt %). In contrast, almandines from S-type magmas (Table 1) and from metapelites have low CaO content (<4 wt %) and variable MnO concentration (Fig. 11a). The compositions of primary almandines of the NPB (Fig. 11a) are consistent with crystallization dominantly from M-type or I-type magmas at high pressure, probably at the mantle–crust boundary zone. Chemical compositions of the amphiboles coexisting with garnets support the high-pressure origin. Most of them crystallized at higher pressures and temperatures than amphiboles in the garnet-free magmas (Fig. 6). On the basis of the CaO and MnO contents of the garnets and the composition of the coexisting amphiboles, Type 1B garnets could have formed at higher pressure than the Type 2 garnets. The mineral assemblage occurring with garnets (Ca-rich plagioclase, magnesiohastingsitic and tschermakitic amphibole) also indicates mantle-derived host magmas. The slightly peraluminous character of the rhyodacites may suggest an S-type origin; however, the relatively high CaO and low MnO contents of the Type 1A garnets (Table 2, Fig. 11a) are inconsistent with an S-type affinity. We suggest that the peraluminous character of the rhyodacites could be the result of contamination by a large amount of lower-crustal metapelitic material. Cores of composite (Type 3) garnets fall in the fields of almandines from S-type magmas and from metapelites (Fig. 11b). Most of them have low MnO contents (<3 wt %), in contrast to the Type 4 almandines from the crustal xenolith (Table 2; Fig. 11b). These garnet fragments may be derived from a heterogeneous lower crust. Type 3 garnet overgrowths have compositions similar to the Type 1B and Type 2 garnets inferred to have a high-pressure origin. This suggests that the xenocrystic Type 3 garnet fragments were incorporated into the host magmas at lower-crustal depths.
Pressure and temperature constraints
Chemical compositions of the NPB garnets and the coexisting minerals suggest that they formed at high pressures and temperatures. Intergrowths of these minerals imply that they are a magmatic equilibrium assemblage and therefore they can be used to quantify the temperature and pressure conditions under which the garnets crystallized. Phase equilibrium diagrams (e.g. Green, 1982, 1992; Huang & Wyllie, 1986; Conrad et al., 1988) may also help to constrain the pressure–temperature conditions of crystallization. The phenocryst assemblage in the NPB volcanic rocks consists of plagioclase, amphibole and/or biotite and garnet. According to experimental results (e.g. Clemens & Wall, 1981; Green, 1982, 1992; Huang & Wyllie, 1986; Conrad et al., 1988), the absence of quartz and presence of plagioclase, amphibole and/or biotite in a dacitic liquid suggest magma H2O content close to 5 wt % and crystallization temperature above 800°C. A higher water content in the magma strongly suppresses plagioclase crystallization, whereas a lower H2O content expands the quartz field and reduces the amphibole field. A mineral assemblage of garnet + plagioclase + amphibole is stable at 8–13 kbar and 820–920°C in a dacitic magma with 5% H2O (Green, 1992).
Among the geothermometers and geobarometers that use the mineral assemblage characteristic of the NPB garnet-bearing volcanic rocks, only two appear to be applicable considering the conditions of calibration. The amphibole–plagioclase (Blundy & Holland, 1990; Holland & Blundy, 1994) and the biotite–garnet (e.g. Ferry & Spear, 1978; Holdaway et al., 1997) thermometers can give the temperature range when garnet crystallized together with these minerals.
Amphibole compositions (high Aliv and Na + K; Fig. 6) in the garnet-bearing volcanic rocks suggest that they crystallized at a relatively high temperature range. Correlation between Aliv and (Na + K)[A] and between Si and Aliv + (Na + K)[A] in amphiboles implies edenite substitution. In addition, Ti-tschermakite substitution could be also important, as indicated by the correlation between Aliv and Ti. Both substitutions are temperature dependent, such that higher temperature can greatly increase Aliv in hornblende (e.g. Helz, 1973; Anderson & Smith, 1995). The Holland & Blundy (1994) amphibole–plagioclase thermometer using a plagioclase compositional range of Xab = 0·2–0·4 and pressure of 7–8 kbar gave an equilibrium temperature of 840–940°C.
Coexisting garnet and biotite may be used to estimate crystallization temperature by applying the Fe–Mg exchange geothermometer originally calibrated by Ferry & Spear (1978). Later calibrations took into account nonideal mixing behaviour in these minerals and incorporated a correction for Mn and Ca content in garnet and Alvi and Ti in biotite (e.g. Indares & Martignole, 1985; Kleeman & Reinhardt, 1994; Holdaway et al., 1997). Holdaway et al. (1997) emphasized also the role of Fe3+ in the calibration compositions. This latter geothermometer yields temperatures from 700°C (Type 1B garnets) to 780°C (Type 1A garnets) for the coexisting garnet and biotite pairs of the NPB volcanic rocks. In contrast, the Ferry & Spear (1978) geothermometer gives significantly higher temperatures, ranging from 730°C (Type 1B garnets) to 880°C (Type 1A garnets). The relatively high temperature for the Type 1A garnets is consistent with the results obtained by Clemens & Wall (1984) for the Violet Town peraluminous rhyolites (830°C) and Barley (1987) for the Mt Somers dacites (880°C) using also the Ferry & Spear (1978) calibration. It indicates an early magmatic equilibrium assemblage, where biotite and garnet are the main near-liquidus ferromagnesian phases (Clemens & Wall, 1981). In contrast, the low temperature (700°C) for the Type 1B garnet and biotite pairs seems to be unreasonably low considering the phase equilibria conditions.
In conclusion, phase equilibria results on dacitic melt and the applied geothermometers indicate that garnets from the NPB volcanic rocks could have formed in the temperature range of 800–940°C. The majority of NPB primary garnets have relatively high grossular content (>10 mol %) consistent with a high-pressure (>7 kbar) origin (Green, 1977, 1992). The absence of quartz and the presence of plagioclase with garnet, amphibole and/or biotite suggest a magma H2O content close to 5 wt %. This volatile content might have enhanced the rapid ascent of the magmas from lower-crustal depths.
Genesis of the garnet-bearing volcanic rocks
Almandine-bearing magmas can be classified as S-type and I- or M-type melts following the scheme for silicic igneous rocks proposed by Chappell & White (1974) and adopted by Green (1992) for garnet-bearing volcanic rocks. Almandine phenocrysts are more common in S-type (peraluminous) magmas (Table 1), i.e. in melts generated by anatexis of metapelitic formations. Such almandines are usually associated with plagioclase, biotite, quartz, K-feldspar and cordierite. Garnet-bearing metapelitic xenoliths are also common in these rocks (e.g. Zeck, 1970; Wood, 1974; Munksgaard, 1984; Barker, 1987). On the other hand, almandine can also occur in I-type magmas, i.e. melts derived by partial melting of igneous rocks in the crust or M-type melts from the upper mantle. These volcanic rocks are diopside-normative and have A/CNK ratios < 1, i.e. metaluminous. In these rocks, almandine has a considerable Ca content (Gro >10 mol %) and coexists with plagioclase, amphibole, biotite, orthopyroxene and rarely with clinopyroxene. Table 1 demonstrates that garnet-bearing I-type and especially M-type volcanic rocks are not as rare as once thought (Green, 1992).
Primary almandine in the NPB volcanic rocks crystallized at 7–12 kbar, i.e. 25–35 km depth. As the present crustal thickness beneath this area is 28–32 km, this depth corresponds to the crust–mantle boundary, as assumed previously by Embey-Isztin et al. (1985) from the experimental results of Green (1982). This also suggests that the primary host magmas were mantle derived. Indeed, the garnet-bearing andesites and some dacites have diopside-normative compositions. On the other hand, the more silicic rocks containing Type 1A garnets are slightly peraluminous and therefore may have been formed by anatexis of metapelites or by extensive contamination of mantle-derived magma. However, garnets in the rhyodacites do not support this origin because their compositions do not resemble almandines in S-type volcanic rocks (Fig. 11a) and they do not coexist with typical ‘S-type’ minerals. Therefore, we conclude that the slightly peraluminous silicic magmas could have evolved from metaluminous liquids either by high-pressure fractionation of clinopyroxene and amphibole (Ellis & Thompson, 1986; Conrad et al., 1988; Green, 1992) or by significant lower-crustal contamination. The isotopic compositions of the host rocks and the narrow curvilinear trend in the Sr–Nd isotope diagram (Fig. 12c) indicate involvement of a crustal component. Source contamination alone is unlikely to be responsible for the increase of radiogenic Sr in the host magmas and the 18O values of garnets (Fig. 12a and b). Assimilation of upper-crustal rocks can be also excluded, because it is not consistent with either the depth (30 km) of origin of the garnets or the Sr–Nd–O isotopic trends. Assimilation of the fusible part of the lower crust could, however, explain these trends (Fig. 12a and b). The low-Ca xenocrystic cores in the Type 3 garnets and their inclusions provide evidence for the presence of metapelitic material in the lower crust beneath the NPB. Embey-Isztin et al. (2000) recently described metasedimentary lower-crustal xenoliths from the Pannonian Basin with a high proportion of Al-rich mineral phases. The linear trend in the 18Ominerals vs (87Sr/86Sr)host rock diagram (Fig. 12a and b) can be explained by two-component mixing between a mantle-derived magma and crustal material with a relatively low 18O value, high 87Sr/86Sr ratio and similar Sr content to that of the magma. The 87Sr/86Sr and 143Nd/144Nd isotope ratios of metasedimentary granulite xenoliths from the Pannonian Basin (Embey-Isztin et al., 2000) and the French Massif Central (Downes et al., 1990; Kempton & Harmon, 1992) are consistent with such a crustal end-member. The high 18O value (10·5) of the Type 3 garnet core from Verce (Börzsöny Mts) indicates that it could be a fragment of this metapelitic lower crust. Mixing calculations suggest 20–25% lower-crustal component in the garnet-bearing andesites and 40–50% lower-crustal component in the rhyodacites. Such a large amount of assimilation requires a high heat flow beneath the NPB during the Miocene. This could be a result of upwelling of hot asthenosphere as a result of the thinning of the lithosphere and the intrusion of mantle-derived magmas.
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Almandine garnets occur only in evolved magmas in the NPB. The strong correlation between most LILE and high field strength elements (HFSE) (e.g. Rb, Ba, La, Th, Zr, Nb) and SiO2 content in the garnet-bearing rocks implies that these elements behaved incompatibly throughout the magma evolution. In contrast, Sr and the HREE decrease in the more silicic rocks, consistent with fractionation of plagioclase, garnet and amphibole. The lack of a strong negative Eu anomaly in the REE pattern of garnet-bearing andesites and dacites (Fig. 10) indicates that plagioclase did not play an important role in the early crystallization stage, but formed simultaneously with garnet. Indeed, the HREE-depleted rhyodacites show a negative Eu anomaly (Eu/Eu* = 0·76–0·87; Fig. 10). The considerable depletion of HREE could be a result of garnet fractionation. Zircon fractionation could also decrease the HREE concentration in evolved liquids, but the incompatible behaviour of Zr in the NPB rocks excludes this process. However, this simple scenario does not explain the anomalously low HREE content of the dacites from the Börzsöny Mts (Fig. 10). These rocks contain <11 ppm Y, whereas the rhyodacites have 10–16 ppm Y, but similar radiogenic isotopic compositions (Table 10). Fractional crystallization could not result in this HREE variation, therefore residual garnet is inferred to have been present during partial melting.
Crystallization of garnet, absence of quartz, and presence of amphibole and/or biotite require some water in the host magma, but the additional presence of plagioclase suggests that the water content cannot be higher than 5 wt % (e.g. Green, 1982, 1992; Huang & Wyllie, 1986). It indicates also a hydrous mantle source unless a significant amount of H2O-rich crustal material was incorporated into the magma. A potential hydrous mantle source could be metasomatized lithospheric mantle. The presence of amphibole and phlogopite in ultramafic xenoliths in alkaline basalts of the NPB (Szabó & Taylor, 1994) and silicate melt inclusions and melt pockets interpreted as infiltration of calc-alkaline melts (Szabó et al., 1996) suggest extensive metasomatism of the lithospheric mantle beneath the NPB. Much of this metasomatism could be related to an earlier subduction event.
Geodynamic implications of eruption of garnet-bearing magmas
Garnet-bearing volcanic rocks are volumetrically rare worldwide. As Mn-poor almandine is not stable at shallow depths (Green, 1977), rapid ascent of a garnet-bearing magma is necessary to preserve this mineral. This requires the presence of deep-seated faults or tectonically weak zones and a tensional stress field. In addition, crystallization of almandine is favoured from silicate liquid with a high water content, so a hydrous mantle source is another essential requirement. These circumstances could be attained in areas that undergo lithospheric stretching following (or concurrent with) subduction. Most garnet-bearing volcanic rocks worldwide appear to be related to such regions. In the Neogene volcanic field of SE Spain (e.g. Doblas & Oyarzun, 1989; Benito et al., 1999; Zeck, 1999), volcanic activity took place during syn- and post-extensional phases, which followed subduction and continental collision. In the Pyrenees, a Permo-Carboniferous garnet-bearing volcanic series formed during the Hercynian orogeny, when extension followed collision (Gilbert & Rogers, 1989). Barley (1987) proposed that the Cretaceous Canterbury volcanic rocks in New Zealand were formed when the tectonic regime changed from compression to extension. The Miocene volcanism of the Northland region (New Zealand) occurred in a post-obduction–subduction tectonic setting, when partial melting took place in a hydrous mantle source (Smith et al., 1989; Day et al., 1992).
In the NPB, the Middle Miocene (13–18 Ma) geodynamic circumstances could also have favoured eruption of garnet-bearing magmas. In the West Carpathians, active subduction was already terminated by this time (e.g. Meulenkamp et al., 1996), but it was the main period of formation of the Pannonian Basin by large-scale lithospheric extension (e.g. Csontos et al., 1992; Tari & Horváth, 1995). The present-day thin crust and especially the thin lithosphere beneath the NPB (Horváth, 1993) imply that this region underwent lithospheric extension. Nemok & Lexa (1990) suggested syn-volcanic extensional tectonic activity and a Basin and Range-type structure for the Central Slovakian Volcanic Field. The garnet-bearing volcanic rocks are remarkably situated along major tectonic lines (Fig. 1b). Therefore, independent lines of evidence support the contention that the Miocene calc-alkaline volcanism in the Northern Pannonian Basin that started with the eruption of garnet-bearing magmas at 16 Ma was related to extension following subduction. The occurrence of garnet-bearing volcanic rocks may imply a change in the tectonic regime in this area and suggest a tensional stress field in Mid-Miocene time.
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ACKNOWLEDGEMENTS |
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Radiogenic and stable isotope and X-ray fluorescence facilities at Royal Holloway are University of London Intercollegiate Research Services. We thank Gerry Ingram and Giz Marriner for their assistance. Andy Beard is acknowledged for helpful assistance during the microprobe work at Birkbeck College. Sz.H.’s research was supported by a NATO post-doctoral research fellowship from the Royal Society. Fieldwork was financially supported by the Hungarian Research Fund (OTKA T025833) and the University of London Central Research Fund. Thanks are due to J. P. Davidson, T. H. Green and J. G. Fitton for their helpful and constructive reviews.
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FOOTNOTES |
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*Corresponding author. Tel: +36-1-266-9833, ext. 2338. Fax: +36-1-266-4992. E-mail: harangi{at}iris.geobio.elte.hu
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REFERENCES |
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