Int J Earth Sci (Geol Rundsch) (2011) 100:107–123 DOI 10.1007/s00531-010-0511-8
ORIGINAL PAPER
Age constraints on Late Paleozoic evolution of continental crust from electron microprobe dating of monazite in the Peloritani Mountains (southern Italy): another example of resetting of monazite ages in high-grade rocks Peter Appel • Rosolino Cirrincione • Patrizia Fiannacca • Antonino Pezzino
Received: 27 October 2008 / Accepted: 18 December 2009 / Published online: 2 February 2010 Ó Springer-Verlag 2010
Abstract In situ monazite microprobe dating has been performed, for the first time, on trondhjemite and amphibolite facies metasediments from the Peloritani Mountains in order to obtain information about the age of metamorphism and intrusive magmatism within this still poorly known sector of the Hercynian Belt. All samples show single-stage monazite growth of Hercynian age. One migmatite and one biotitic paragneiss yielded monazite ages of 311 ± 4 and 298 ± 6 Ma, respectively. These ages fit with previous age determinations in similar rocks from southern Calabria, indicating a thermal metamorphic peak at about 300 Ma, at the same time as widespread granitoid magmatism. The older of the two ages might represent a slightly earlier event, possibly associated with the emplacement of an adjacent trondhjemite pluton, previously dated by SHRIMP at 314 Ma. No evidence for preHercynian events and only a little indication for some monazite crystallization starting from ca. 360 Ma were obtained from monazite dating of the metasediments, suggesting either a single-stage metamorphic evolution or a significant resetting of the monazite isotope system during the main Hercynian event (ca. 300 Ma). Rare monazite from a trondhjemite sample yields evidence for a lateHercynian age of about 275 Ma. This age is interpreted as representing a post-magmatic stage of metasomatic monazite crystallization, which significantly postdates the emplacement of the original magmatic body. P. Appel Institut fu¨r Geowissenschaften, Universita¨t Kiel, 24098 Kiel, Germany R. Cirrincione P. Fiannacca (&) A. Pezzino Dipartimento di Scienze Geologiche, Catania University, Corso Italia 57, 95129 Catania, Italy e-mail:
[email protected]
Keywords Peloritani Mountains Sicily Monazite microprobe dating Metapelite Trondhjemite
Introduction In the last decades the crustal evolution of the Calabria– Peloritani Orogen has been the subject of considerable work and Hercynian metamorphism and late-Hercynian magmatism of both southern and northern sectors of the orogen have started to be constrained on geochronological grounds (Schenk 1980, 1990; Graeßner et al. 2000). Nevertheless, for the southernmost termination of the Calabria– Peloritani Orogen (Peloritani Mountains, north-eastern Sicily), ages for magmatism and/or metamorphism are still very scarce and mostly based on Rb/Sr and Ar–Ar systems. A single high-precision in situ U–Pb zircon age of 314 ± 3 Ma was recently obtained for a trondhjemitic sample (Fiannacca et al. 2008). A complicated structure and a polyphase tectono-metamorphic history of the Peloritani Mountains, involving Alpine, Hercynian and possibly pre-Hercynian events is postulated by most authors studying this area (e.g. Atzori and Ferla 1992; Cirrincione and Pezzino 1994; Ferla 2000 and references therein). In such a context, in situ chemical dating of monazite appears as a complementary method to obtain reliable constraints useful to unravel the pre-Alpine history of the Peloritani Mountains. Electron microprobe dating of monazite is a non-destructive method, which offers the possibility of dating monazite in its textural context allowing correlations to be made between age data and discrete magmatic and subsolidus events. The Aspromonte–Peloritani Unit, cropping out in the north-eastern part of the Peloritani Mountains and in the Aspromonte Massif of southern Calabria (Fig. 1) was selected for this study since it stands out for its
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Fig. 1 Geological sketch map of Peloritani Mountains with location of studied samples (dark rectangles). a Structural sketch map of the Peloritani Mountains and the Aspromonte Massif. b Basement-granitoids relationships in the Calabria– Peloritani Orogen. c Detailed field map
widespread occurrence of trondhjemite plutons, which are lacking in the other tectonic units of the Calabria Peloritani Orogen. The present study reports the first monazite ages obtained for two samples of high-grade metasedimentary rocks and for a trondhjemite sample cropping out in the Aspromonte–Peloritani Unit of the north-eastern Peloritani. It is the aim of this study to contribute to a better understanding of the link between the tectono-metamorphic evolution, including melt-generation and the post-magmatic processes, in this segment of the southern European Hercynian Belt.
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Geological setting The Peloritani Mountain Belt (north-eastern Sicily) constitutes the south-western termination of the Calabria– Peloritani Orogen, a remnant of the Hercynian Belt, which was reworked during the Alpine orogeny and that now connects the Southern Apennine Chain and the Maghrebid Chain (Fig. 1). The Peloritani Mountains consist of a set of south-verging nappes of Hercynian basement rocks, with metamorphic grade increasing towards the top, and interposed fragments of Mesozoic–Cenozoic
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sedimentary covers (Atzori and Vezzani 1974; Lentini and Vezzani 1975). The entire Peloritani Belt may be subdivided into two complexes with different tectono-metamorphic histories (Atzori et al. 1994; Cirrincione and Pezzino 1994). The Lower Domain is exposed in the southern part of the Peloritani Belt and comprises volcano-sedimentary Cambrian–Carboniferous sequences, which were affected by Hercynian sub-greenschist to greenschist facies metamorphism and which now are covered by unmetamorphosed Mesozoic–Cenozoic sediments. The Upper Domain, in the north-eastern part of the belt, consists of two units (Mandanici Unit and Aspromonte–Peloritani Unit), showing Hercynian greenschist to amphibolite facies metamorphism, which in part are also affected by a younger Alpine greenschist facies metamorphic overprint (Cirrincione and Pezzino 1991; Atzori et al. 1994). Fragments of a metamorphosed Mesozoic–Cenozoic cover occur interposed between the Mandanici Unit and the Aspromonte–Peloritani Unit. In the Peloritani Belt the Aspromonte–Peloritani Unit represents the highest tectonic Unit, overlying the Mandanici Unit, whereas in the Aspromonte Massif of southern Calabria the same unit is sandwiched between the lowermost Madonna di Polsi Unit (Pezzino et al. 2008) and the uppermost Stilo Unit (Fig. 1). The most prominent rock types in the Aspromonte–Peloritani Unit, in both northern Sicily and southern Calabria, are middle crustal biotite paragneisses and augen gneisses with minor amphibolites, mica schists and marbles. The metamorphic rocks are diffusely intruded by late-Hercynian weakly to strongly peraluminous granitoids (D’Amico et al. 1982; Rottura et al. 1990, 1993; Fiannacca et al. 2005a, 2008). Late-Hercynian granitoids belong to two different suites: a main (representing ca. 70% of the exposed rocks) metaluminous to weakly peraluminous calc-alkaline batholitic suite, and a strongly peraluminous suite, which is composed of a number of small scale intrusions (Rottura et al. 1993 and references therein). The granitoids are lateto post-tectonic and were probably emplaced along ductile shear zones connected to an extensional regime (Rottura et al. 1990; Caggianelli et al. 2000, 2007). Only relatively small plutons, mostly composed of weakly to strongly peraluminous granitoids, are exposed within the mediumhigh-grade basement of the Aspromonte–Peloritani Unit. Among these granitoids, weakly peraluminous trondhjemites, not included in the frame of the late-Hercynian magmatism by most of the previous authors, crop out either as small plutons and stocks and as dm to m discordant to sub-concordant leucosomes and dykes. Preliminary studies on some of the trondhjemite bodies have suggested that their origin is related to Hercynian crustal evolution processes (Atzori et al. 1984a; Lo Giudice et al. 1985;
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Fiannacca et al. 2005a). An origin by alkali metasomatism at the expense of late-Hercynian granitoids was assumed by Fiannacca et al. (2005a) for the Pizzo Bottino trondhjemites, and a U–Pb zircon SHRIMP age of 314 ± 3 Ma has been obtained for the emplacement of the original magmatic body (Fiannacca et al. 2008).
Metamorphic conditions and previous geochronological work The timing of formation of the crystalline basement of the Peloritani Mountains is still an open question. Most authors ascribed it entirely to the Hercynian orogeny (e.g. Atzori et al. 1984b; Ioppolo and Puglisi 1989; Messina et al. 1996), whereas others suggested that the Mandanici and Aspromonte Units represent part of a pre-Hercynian polymetamorphic basement (e.g. Ferla 1978, 2000; Bouillin et al. 1987) or the result of welding of different preHercynian and Hercynian terranes during the final stages of Hercynian orogeny (De Gregorio et al. 2003). P–T estimates for the rocks of the Aspromonte–Peloritani Unit in the Peloritani Mountains range between ca. 680 and 550°C at ca 5.0–3.0 kbar (Atzori et al. 1984b; Ioppolo and Puglisi 1989; Messina et al. 1996; Rotolo and De Fazio 2001), with P–T peak conditions similar to those estimated for the rocks of the Central Aspromonte Massif (650– 675°C at 4.0–5.0 kbar; Ortolano et al. 2005). Festa et al. (2004, and references therein) additionally reported the occurrence of pre/eo-Hercynian mafic granulites within Hercynian migmatites of the northern Peloritani, with Ca-rich Grt ? Cpx ? Qtz assemblages suggesting P & 8–10 kbar and T & 700°C. A final widespread episode of hydration under decreasing temperatures (&480°C) was probably caused by the massive emplacement of metaluminous to strongly peraluminous late-Hercynian granitoids at about 290 Ma (Rb–Sr data on micas; Rottura et al. 1990). As for the pre-Hercynian evolution of southern Calabria–Peloritani Orogen, different authors (Schenk and Todt 1989; Schenk 1990; Micheletti et al. 2007; Fiannacca et al. 2008, 2009) reported U–Pb zircon data from different levels of the southern Calabrian and north-eastern Sicilian crust indicating a late Pan-African/Cadomian (600– 500 Ma) crust-forming event. Micheletti et al. (2007) obtained SIMS ages in the range of c. 560–525 Ma for the magmatic protoliths of Calabrian augen gneisses, with Archean (3.1 Ga), Paleoproterozoic (1.7–2.4 Ga) and Neoproterozoic (0.6–0.9 Ga) inheritance. A similar set of inherited ages and additional Pan-African inheritance has been revealed by SHRIMP dating of zircon from a lateHercynian leucogranodiorite (Fiannacca et al. 2008). De Gregorio et al. (2003) obtained a wide range of Ar–Ar
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(stepwise heating method), Rb–Sr and U–Pb (stepwise leaching method) ages for amphibole, biotite, muscovite, K-feldspar and titanite, for samples of different rock types from the north-eastern Peloritani, starting from about 1.6– 1.8 Ga up to 61–48 Ma. The latter ages are interpreted to be related to Tertiary events localized along narrow shear zones. The only pre-Hercynian ages are Ar–Ar hornblende ages, from two amphibolite samples, interpreted as mixing ages between a younger generation formed around 600 Ma with older cores; U–Pb dating of one titanite from one of those samples gave a still older age of 1.6–1.8 Ga. Geochronological indications for early-Hercynian events have been reported, from metapelites of southern Calabria, by a Rb/Sr biotite age of ca. 330 Ma (Bonardi et al. 1987), and by a poorly constrained lower concordia intercept age of 377 ± 55 Ma (Schenk 1990). Nevertheless, more recently, Bonardi et al. (2008) indicated for the same rocks Rb/Sr muscovite ages of ca. 314–308 Ma, which are closer to recently obtained ages for the Hercynian metamorphism. In Sicily, De Gregorio et al. (2003) reported eo-Hercynian 39Ar–40Ar hornblende ages ranging from 420 to 350 Ma and Hercynian 39Ar–40Ar hornblende and muscovite ages in the range of 340–300 Ma. Nevertheless, the above authors interpret the first group of ages as the ages of pre-Hercynian magmatic cumulates that escaped complete Hercynian resetting, whereas no interpretation is provided for most of the ages of the second group. Ar–Ar amphibole ages of ca. 300 Ma were interpreted as the age of retrograde metamorphism of a former cpx–grt peak assemblage, while a muscovite age of 301 ± 2 Ma from mylonitic augen gneiss was interpreted to be related to nappe stacking. The same authors obtained Ar–Ar biotite ages roughly in the range ±240–50 Ma for the biotitic paragneisses from the northern Peloritani. They reach the conclusion that the basement of north-eastern Peloritani was mainly built in the Paleozoic, but that pre-Hercynian relics and localized Tertiary overprints also occur. Atzori et al. (1990) indicated a common metamorphic history for augen gneisses and associated biotitic paragneisses from the north-eastern Peloritani with Rb/Sr ages on micas of 280–292 Ma, interpreted as cooling ages after the Hercynian metamorphism. U–Pb monazite ages (Graeßner et al. 2000) for similar amphibolite facies paragneisses of the Aspromonte Massif indicated a metamorphic peak at 295 to 293 ± 4 Ma (with P–T conditions of 620°C at ca. 2.5 kbar for the base of the upper crust), coeval with the lower crust (exposed in the Serre Massif, southern Calabria), the latter characterized by a peak temperature of 690–800°C at 5.5–7.5 kbar (Graeßner et al. 2000 and reference therein). This metamorphic peak was nearly synchronous with the bulk of granitoid intrusions at 303–290 Ma (zircon, monazite and xenotime
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U–Pb ages and whole-rock and mineral Rb–Sr ages; Borsi and Dubois 1968; Borsi et al. 1976, Schenk 1980; Del Moro et al. 1982; Graeßner et al. 2000; Fiannacca et al. 2008). According to Graeßner and Schenk (1999) and Graeßner et al. (2000) emplacement and crystallization of the large volumes of granitoids were possibly responsible for a regional scale late-stage metamorphism. This event occurred under static conditions and resulted in a complete recrystallisation of the mineral assemblages thus erasing almost all records of previous tectonic and magmatic/ metamorphic stages, which are now only rarely preserved (Caggianelli et al. 2007 and references therein). Despite the absence of detailed geochronological constraints, clockwise P–T–(t) paths inferred for the medium- to high-grade rocks of the Aspromonte Massif and Peloritani Mountains have been considered to be consistent with processes of crustal thickening during early- and middle-Hercynian collisional stages, followed by crustal thinning, granitoid intrusion and unroofing during lateHercynian extensional stages (Festa et al. 2004; Caggianelli et al. 2007). In particular, Atzori and Ferla (1992) proposed distinct peaks for P and T in the northern Peloritani Mountains, with eo-Hercynian, or pre-Hercynian, P peak (at ca. 600°C and 5.5 kbar) predating late-Hercynian thermal peak (at ca. 630°C and 4.0 kbar). Data reported for the Peloritani Mountains appear to be consistent with a stage of low-P/high-T metamorphism at about 300 Ma overprinting an older phase of Barrovian metamorphism, as suggested for the whole orogenic segment (Graeßner and Schenk 1999; Graeßner et al. 2000; Caggianelli and Prosser 2002; Festa et al. 2004).
Field relations and petrography Amphibolite facies paragneisses and migmatites represent the most common rock types in the Aspromonte–Peloritani Unit. These rocks are the result of Hercynian metamorphism of greywacke sequences with variable pelitic components (Lo Giudice et al. 1985, 1988). Their typical mineral assemblage consists of quartz–plagioclase–biotite with variable amounts of muscovite, K-feldspar, sillimanite, garnet, cordierite and andalusite. Synkinematic crystallization appears to dominate in most paragneisses, but a strong textural reorganization associated with crystallization of plagioclase and micas is commonly evident, testifying for post-deformational events. Migmatites are mostly metatexites, with main foliation defined by alternating millimetre- to decimetre-scale leucosomes and biotite-rich meso-melanosomes. Leucosome composition in the Peloritanian migmatites is leucogranitic to trondhjemitic (Maccarrone et al. 1978; Atzori et al.
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1985; Fiannacca et al. 2005b). Leucosomes studied by Fiannacca et al. (2005b) have a trondhjemitic composition, similar to melts produced by H2O-fluxed melting of muscovitic schists (Patin˜o Douce and Harris 1998). However, from petrographic investigations such trondhjemitic compositions often result to be a kind of ‘‘average’’ composition, which derives from the mixing of different compositional domains. Similar ‘‘mixed’’ leucosomes, suggesting multi-stage melting events, are reported for the Serre metapelitic migmatites (Fornelli et al. 2002). The mineralogical composition of Peloritanian leucosomes appears further complicated by the presence of little amounts of restite phases and the later occurrence of melt– solid phases back-reactions and subsolidus alteration (Fiannacca et al. 2005b). Trondhjemitic rocks constitute variously sized scattered bodies displaying a variety of field relations with the surrounding metamorphic rocks (Fiannacca et al. 2005a and references therein). The largest bodies are about 10 km2 in extension and their contacts with the country rocks vary from sharp (either discordant or concordant) to gradual. Trondhjemites are sometimes associated with other leucocratic rocks of granodioritic–monzogranitic compositions, which have the same appearance in the field. Trondhjemites from the Peloritani Mountains are leucocratic and mostly coarse- to very coarse-grained heterogranular rocks. Variously sized metasedimentary enclaves of restitic/ xenolithic significance are frequently observed. The texture is hypidiomorphic to autoallotriomorphic but a strong subsolidus crystallization, related to metasomatism and retrogression coupled with common solid-state textural readjustment at depth, overprints the original magmatic features, which are preserved within domains of variable size.
Samples Sample GC13 was collected close to a small trondhjemitic body (Fig. 1). It is a fine-grained paragneiss composed of quartz, plagioclase, biotite and very scarce K-feldspar. Accessory phases are ilmenite, zircon, monazite and unusually abundant apatite. Muscovite and chlorite only occur as retrograde phases, grown at the expense of previously formed biotite. The sample displays a main foliation S1 marked by synkinematic growth of biotite (Bt1), plagioclase (Pl1), quartz, ilmenite and apatite. The S1 schistosity is partly affected by a coarse crenulation, which does not lead to the formation of a new schistosity. A second generation of coarser biotite (Bt2) cuts randomly the foliation suggesting postkinematic growth. Thermobarometric data for similar paragneisses from the same area suggest similar T conditions for the synkinematic and the
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postkinematic stages, developed at T of ca. 550 and 550– 500°C, respectively. P conditions, estimated on the basis of the phengite content of muscovite coexisting with K-feldspar (Massonne and Schreyer 1987) and of the Ca distribution between garnet–plagioclase pairs (Ghent et al. 1979, range from ca. 3.6 kbar, estimated for the synkinematic stage, to 3.4–3.0 kbar obtained for the postkinematic one (Ioppolo and Puglisi 1989). Sample PB14 is from a metapelitic migmatite cropping out at the southern side of the Pizzo Bottino trondhjemite body (Fig. 1), which represents one of the largest trondhjemite occurrences in the Peloritanian area (Fiannacca et al. 2005a, b). The leucosomes (millimetre- to decimetresized) consist of the typical mineral assemblage quartz– plagioclase–muscovite–K-feldspar–sillimanite with minor biotite and apatite and rare zircon and monazite as accessories. They are always concordant within the foliation and occur as lenticular bodies and as discontinuous layers, which are often rimmed by millimetre-sized melanosomes. Grain size of leucosomes is mainly coarse to very coarse. PB14 leucosome-forming minerals are quartz, plagioclase, muscovite, K-feldspar, sillimanite and small amounts of biotite; accessory phases are apatite, zircon and rare monazite. The overall composition of the leucosomes is trondhjemitic with K-feldspar occurring as scattered inclusions in plagioclase crystals or, less frequently, as an interstitial and subhedral phase. Evidence for muscovite breakdown under anhydrous conditions according to the reaction muscovite ? quartz = K-feldspar ? sillimanite ? melt is locally given by textural relationships and point to temperatures in excess of ca. 650°C, at pressure greater than 3.5 kbar, as inferred for muscovite dehydration melting in the stability field of sillimanite. Widespread growth of secondary plagioclase plus myrmekites and muscovite–quartz symplectites at the expense of magmatic microcline testifies the occurrence of significant crystallization following the anatectic stage (Fiannacca et al. 2005b). Meso and melanosomes show medium to fine grain sizes and grano-lepidoblastic textures with some portions characterized by postkinematic blastesis of micas and plagioclase and the development of a granoblastic polygonal texture. Mesosomes forming minerals are quartz, plagioclase, biotite, muscovite and very little amounts of K-feldspar. Accessory phases are apatite, zircon, monazite, Fe–Ti oxides. Biotite is synkinematic, in single plats or in lepidoblastic associations or, more rarely, postkinematic in single plats. Muscovite is mostly synkinematic. Melanosomes consist of millimetre-sized mafic selvages rimming leucosomes and are mainly composed of biotite with limited amounts of sillimanite and muscovite. T estimates available for similar Peloritanian migmatites (biotite–garnet and two-feldspar geothermometry; Ioppolo and Puglisi 1989) indicate T of ca. 640°C and 650–600°C for the
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synkinematic and the postkinematic stages, respectively; the postkinematic stage appears to have evolved at decreasing pressures, in the range of 4.0–3.2 kbar. Sample GC30 is a sample of a trondhjemite collected from the Pizzo Bottino rock body (Fiannacca et al. 2005a; Fig. 1). The selected sample is mainly composed of plagioclase and quartz (ca. 90 vol%), and contains small amounts of biotite, muscovite and K-feldspar. Accessory phases are apatite, zircon, monazite and Fe–Ti oxides. The sample is characterized by the occurrence of different plagioclase populations represented by prevailing anhedral to subhedral megacrysts and less abundant millimetre-sized crystals. Several of the latter plagioclase crystals display magmatic features, such as euhedral elongated habit and simple twinning. Quartz forms medium to large anhedral discrete grains or glomerocrystic aggregates; anhedral or rounded quartz also occurs within the plagioclase. Biotite and muscovite mainly occur as euhedral plates of variable size with frequently corroded or fringed rims within the plagioclase. Microcline occurs as scattered inclusions in large plagioclase and in very rare interstitial patches.
Monazite age dating Analytical techniques All monazite analyses were performed with a JEOL JXA 8900 microprobe at the University of Kiel, equipped with five WD spectrometers. For each analytical point a full monazite analysis, consisting of 15 elements, was performed. An accelerating voltage of 20 kV, a probe current of 80 nA and a focused beam were used for all monazite analyses. The JEOL H-type spectrometer with reduced Rowland circle for high count rates was used for measurements of lead, thorium and uranium. Background offsets were selected after long time fine WD scans of natural monazite. The interference of Th Mc on U Mb was corrected with an experimentally determined correction factor. Counting times were adapted to net intensities to achieve the desired objective of a low error for counting statistics at reasonable counting times. Typical counting times for Pb are 240 s on the peak and background. Under these conditions and using natural crocoite as calibration standard, the theoretical 1-sigma detection limit for Pb, as calculated from the counting statistics of the background, is in the range of 40–60 ppm. As standard materials, synthetic REE orthophosphates (Jarosewich and Boatner 1991) corrected for their Pb contents (Donovan et al. 2003) were used for P, REE and Y; synthetic U-bearing glass for U; natural wollastonite for Ca and Si; natural thorianite for Th; natural
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crocoite for Pb; and corundum for Al. The JEOL ZAF program was used for matrix correction of the monazite analyses. All monazite grains were measured in situ in petrographic thin sections. A summary of the electron microprobe data for the analysed samples is given in Table 1. Special polishing methods were used to avoid contamination by lead during sample preparation. To control the quality of the data the internal laboratory standard F6 from Manangoutry Pass, SE Madagascar (kindly provided by Michael Raith, Bonn), was repetitively analysed during the measurements. This monazite is homogeneous in composition and was dated with the U–Pb method of cogenetic zircon and by a Sm–Nd monazite–biotite–garnet–zircon isochron at 545 ± 2 and 542 ± 11 Ma, respectively (Paquette et al. 1994), and recently by A. Mo¨ller (pers. comm.) with TIMS at an age of 560 ± 1 Ma. Data with significant Al contents were rejected to eliminate analyses with secondary fluorescence artefacts that may occur if analytical points are close to grain boundaries. The analytical error for Th, U and Pb for each point was calculated from counting statistics of the unknown and the standard, using own software (for download at http://www.ifg.uni-kiel.de/213.html). After each monazite measurement, the apparent age was calculated by solving the decay equation for the U–Th–Pb system iteratively. All calculated apparent ages of each sample were then evaluated to identify analyses that clearly fall outside the range defined by a bell-shaped normal distribution. This approach, together with a check of the petrographic position of each measurement point, helps to identify mixing ages without geological significance. Two methods were used to calculate ages for a set of data: (1) the widely used chemical isochron method (CHIME, Suzuki et al. 1991) and (2) the calculation of a weighted average, which is frequently used in more recent studies on chemical dating of monazite (e.g. Pyle et al. 2005). Isochron calculations were performed with the CHIME-program of Kato et al. (1999). The isochron plot has the advantage that it easily enables the identification of data sets with distinct ages and additionally gives information about the chemical variation of the analysed monazite in terms of ThO2* and PbO. A drawback of this method is that it typically suffers from high errors, especially in cases when the variation of ThO2* is small, and thus leads to a poorly defined regression line. In contrast to this, the error of the weighted mean does not depend on the chemical variation of the data and typically yields errors that are significantly lower than the error calculated from the isochron method.
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Table 1 Representative electron microprobe analysis and apparent ages of monazite from studied samples Sample no.
GC13 82
PB14 102
117
61
GC30 75
94
32
33
35
P2O5
31.10
30.91
30.65
30.71
26.39
30.82
30.66
30.58
30.83
SiO2
0.24
0.33
0.32
0.13
2.23
0.07
1.00
0.31
0.24
CaO
0.64
0.81
0.96
0.92
2.05
1.17
3.11
3.13
3.21
La2O3
15.28
14.36
13.83
14.44
10.22
12.93
8.21
8.30
8.05
Ce2O3
28.98
28.27
27.76
28.99
21.63
27.56
20.42
20.69
20.10
Pr2O3
3.09
3.02
3.05
2.97
2.41
2.98
2.45
2.47
2.52
Nd2O3 Sm2O3
11.91 1.90
12.17 2.10
11.89 1.99
11.39 1.69
9.24 1.71
11.67 2.05
9.19 3.52
9.26 3.60
9.35 3.65
Eu2O3
0.00
0.00
0.00
0.00
0.00
0.00
0.00
0.15
0.21
Gd2O3
1.61
1.67
1.69
1.38
1.40
1.70
2.68
2.81
2.95
Dy2O3
0.51
0.49
0.60
0.41
0.64
0.64
1.03
1.06
1.15
Er2O3
0.13
0.05
0.06
0.15
0.15
0.14
0.08
0.08
0.10
Y2O3
1.97
1.90
2.04
2.18
2.17
2.39
2.79
2.98
3.07
ThO2
3.04
4.24
4.78
3.86
17.70
4.32
5.25
5.10
5.23
UO2
0.40
0.28
0.60
0.60
0.84
1.17
9.40
9.78
9.76
PbO
0.06
0.06
0.08
0.07
0.29
0.10
0.43
0.44
0.43
Al2O3
0.00
0.00
0.00
0.00
0.00
0.00
0.43
0.05
0.02
100.86
100.66
100.31
99.90
99.07
99.68
100.63
100.77
100.85
Total P
1.006
1.003
1.000
1.005
0.907
1.009
0.983
0.995
1.000
Si
0.009
0.013
0.012
0.005
0.090
0.003
0.038
0.012
0.009
Ca
0.026
0.033
0.040
0.038
0.089
0.049
0.126
0.129
0.132
La Ce
0.215 0.405
0.203 0.397
0.197 0.392
0.206 0.410
0.153 0.322
0.184 0.390
0.115 0.283
0.118 0.291
0.114 0.282
Pr
0.043
0.042
0.043
0.042
0.036
0.042
0.034
0.035
0.035
Nd
0.163
0.167
0.164
0.157
0.134
0.161
0.124
0.127
0.128
Sm
0.025
0.028
0.027
0.023
0.024
0.027
0.046
0.048
0.048
Eu
0.000
0.000
0.000
0.000
0.000
0.000
0.000
0.002
0.003
Gd
0.020
0.021
0.022
0.018
0.019
0.022
0.034
0.036
0.037
Dy
0.006
0.006
0.007
0.005
0.008
0.008
0.013
0.013
0.014
Er
0.002
0.001
0.001
0.002
0.002
0.002
0.001
0.001
0.001
Y
0.040
0.039
0.042
0.045
0.047
0.049
0.056
0.061
0.063
Th
0.027
0.037
0.042
0.034
0.164
0.038
0.045
0.045
0.046
U
0.003
0.002
0.005
0.005
0.008
0.010
0.079
0.084
0.083
Pb
0.001
0.001
0.001
0.001
0.003
0.001
0.004
0.005
0.005
Al
0.000
0.000
0.000
0.000
0.000
0.000
0.019
0.002
0.001
Total
1.992
1.992
1.994
1.995
2.005
1.994
2.001
2.001
2.000
Xhu
0.004
0.007
0.008
0.002
0.085
0.001
0.003
0.004
0.001
Xmnz Xxe
0.872 0.070
0.857 0.068
0.838 0.073
0.850 0.071
0.663 0.076
0.819 0.082
0.627 0.108
0.624 0.112
0.615 0.117
Xch
0.054
0.068
0.081
0.078
0.177
0.099
0.263
0.260
0.267
App. age (Ma)
309.4
Err (2-sigma)
46.3
271 38.4
274.6
298
338
300
284.8
280.9
277.5
30.2
36
12
25
9.1
8.5
8.1
Cations calculated on the basis of 4 oxygen
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Chemical composition and ages of monazite Paragneiss GC13 Monazite of sample GC13 is generally smaller than 100 lm and anhedral with irregular-shaped grains (Fig. 2a–c). Almost all grains are rich in small rounded inclusions and have cracks, along which alteration has affected the monazite. Some monazites are dismembered in their outer parts. No significant zonation was observed in backscattered electron images (BSEI). Monazite in this sample either occurs in close association with biotite or within the matrix. The first type of monazite is grown along the grain boundaries of biotite or within biotite along cracks. Matrix monazites are mostly smaller and more rare. The chemical variation of monazite of GC 13 is relatively restricted. According to the nomenclature of Linthout (2007) all plot within the compositional field of monazite (Fig. 3a). Th and LREE contents vary systematically, but to a small extent (Fig. 6a). U and Th do not show any correlation (Fig. 4) and the plot of Y versus HREE shows an unsystematic scatter of data points in a wide range of the diagram (Fig. 6b). In the Th ? U ? Si versus REE ? Y ? P diagram (Fig. 3b) the analyses plot along a straight line close to the cheralite exchange vector, thus indicating a single-stage growth of monazite and the incorporation of Th and U into the monazite structure mainly due to the cheralite exchange. Fig. 3 a Nomenclature diagram for the system 2REEPO4– CaTh(PO4)2–2ThSiO4 after Linthout (2007). Analyses of all investigated samples fall within the compositional range of normal monazite. Monazite of sample GC30 has a higher cheralite component and also some patchy domains in monazite of sample PB14 have significant huttonite and cheralite component. b Plot of REE ? Y ? P versus Th ? U ? Si concentrations of monazites from studied samples
Fig. 2 Backscattered electron images of representative monazite grains in paragneiss (a–c sample GC13), migmatite (d–f sample PB14) and trondhjemite (g–i sample GC30) from the Aspromonte– Peloritani Unit of north-eastern Peloritani
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The individual ages of these analyses scatter in a wide range from 143 to 361 Ma. Plotted in a histogram (Fig. 7a), they show a bimodal distribution with one maximum at about 295 Ma and a second maximum at approximately 220 Ma. To calculate an age for the early phase of monazite growth, analyses with apparent ages younger than 250 Ma cannot clearly be assigned to the younger or older Gauss curve. All analyses with ages less than 250 Ma were therefore excluded from the calculations of the high age. The isochrone of sample GC13 thus is based on 38 analyses and yields an age of 302 ± 58 Ma (2-sigma) (Fig. 8a). The weighted mean age of 298 ± 6 (2-sigma) Ma for these analyses is almost identical to the isochrone age. The reason for the large scatter with a bimodal age distribution of the individual ages for sample GC13 first
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(a) 0.04 GC13 PB14 GC30
0.03 0.02
Δ PbO (wt.-%)
0.01 0 -0.01 -0.02 -0.03
(b)
-0.04
0.06
appears to be unclear. There is no general pattern within the age data which corresponds to the textural position of the analyses. Moreover, within individual grains the ages scatter but most of the younger ages occur in the more rimward parts of the grains. To deduce the significance of the younger ages of this sample, we calculated in a first step a theoretical PbO content for each data point, which is calculated from the weighted mean age for GC13 (298 Ma) and from the measured ThO2 and UO2 concentrations. If the Pb in this sample is undisturbed and not affected by recrystallization and/or diffusion, the difference of this value and the measured PbO has to correlate with the difference between the calculated age of the analyses and the mean age of the sample (Fig. 5a). Further, if a second stage of monazite growth younger than that of the used mean age (298 Ma) had occurred, then the respective data points should be shifted towards DPbO values above the linear data arrangement that is defined by data that reflect an age of 298 Ma. Figure 5a shows a very good correlation for most data points, and only few data points with low apparent ages plot above the linear data arrangement. Y contents of sample GC13 scatter in a wider range than in the other samples and show only a very weak dependence on the age (Fig. 5b). The unsystematic behaviour of Y and HREE suggests that monazite in this sample underwent disturbance in the distribution of these elements after crystallization of the monazite to different degrees. This, however, cannot be applied to Pb, Th and U because the systematic correlation
0.05
Y (c.p.f.u.)
Fig. 4 Plot of UO2 versus ThO2 concentrations. No systematic correlation between these elements can be observed. Sample GC30 contains monazite with high U-rich content and sample PB14 shows a wide range of ThO2 concentrations
0.04
0.03
0.02 0.01 -300
-200
-100
0
100
200
300
Δ Age (Ma) Fig. 5 a DPbO versus Dchemical age diagram. DPbO is the difference between a hypothetical PbO content, calculated from the measured ThO2, UO2 and the weighted mean age of the sample and the measured amount of PbO. Dchemical age is simply the difference between the apparent age calculated for each analyses and the weighted mean age for the sample. In case that the element distribution of PbO, ThO2 and UO2 is undisturbed and the used mean age is correct, all data points should plot along a straight line, which crosses the origin of the axis. In case that a younger population of data exists these data points should be shifted to higher DPbO. b Chemical variation of Y (cations per 4 oxygen) versus Dchemical age. The plot displays a weak tendency for low apparent ages of analyses with low Y
of PbO to the reference of the mean age of 298 Ma as displayed in Fig. 5a indicates that these elements are not significantly affected by such element re-distribution. Furthermore, the overall linear data arrangement in Fig. 5a invalidates the interpretation that the young ages in the age histogram (Fig. 7a) are geologically significant. It is thus believed that monazite in GC13 crystallized during a single-stage event at ca. 298 Ma.
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Migmatite PB14 The metapelitic sample PB14 contains abundant monazite. Most of it occurs within the meso-melanosome, whereas monazite in the leucosome is very rare. Monazite is closely intergrown with biotite or can be found in the matrix close to the grain boundaries with biotite. The grains are variable in size, ranging from subhedral larger ones (50–100 lm) to small anhedral grains which are mostly rounded (Fig. 2d– f). Monazite is unzoned in BSEI except for some grains, which include irregular-shaped bright patches that are about 10–20 lm in size. These domains are enriched in cheralite (up to 25 mol%) and huttonite (up to 17 mol%) component with ThO2 contents up to 27 wt% and low Ce2O3 values. Except from these, all analyses fall in the compositional range of monazite (Fig. 3a) and have a relatively restricted range of ThO2 between 3.5 and 6 wt% (Fig. 4). Due to the overall large variation in ThO2 also the LREE vary systematically, whereas Y and the HREE only cover a restricted compositional range (Fig. 6a, b). The apparent ages obtained from the analyses also vary in a restricted range; 51 of a total of 60 analyses yield apparent ages that are inside the range of a normal distribution between 290 and 340 Ma. The isochrone age, based on this set of analyses is 306 ± 20 Ma (2-sigma) and the weighted mean age is 311 ± 4 (2-sigma) Ma (Figs. 7b, 8b). Again both ages are very close to each other. Also, the ages of the Th-rich patches fall within the range of the weighted mean age. Trondhjemite GC 30 Only few grains of monazite were found in sample GC30. They generally exhibit elongated shapes and are unzoned
Fig. 6 Chemical variation diagrams for monazite analyses. a LREE versus Th. These elements show a good correlation, which is mainly caused by the huttonite exchange Th4? ? Si4? = LREE3? ? P5?. b HREE versus Y plot. Y substitutes for REE and behaves similar to
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(Fig. 2g–i). Most grains are very small (\10 lm) and therefore only very few analyses could be obtained. All show that monazite is enriched in cheralite component with low huttonite component (Fig. 3a), and high content of UO2 of more than 9 wt% (Fig. 4). Also, monazite from this sample has the highest contents of Y and HREE which are positively correlated (Fig. 6b). LREE contents of this sample are relatively low (ca. 0.6 per 4 oxygen, Fig. 6a). The ages of these analyses range between 250 and 290 Ma, when calculated from ThO2*, but age calculation from UO2* yields similar results. The weighted mean age is 275 ± 4 Ma (2-sigma). Because only five analyses could be obtained from two tiny grains, no isochron was calculated for this sample. The ages represent only a rough and provisional estimate for the time of monazite crystallization and should be interpreted with caution. Nevertheless, at the current state of knowledge they point to a somewhat younger event compared to the ages of the semipelite and the metapelite.
Discussion Chemical dating with the electron microprobe was done on monazite of two amphibolite facies metasedimentary rocks and one trondhjemite from the basement of the Peloritani Mountains. The data obtained by monazite dating of paragneiss and migmatite samples suggests a Hercynian evolution for both lithotypes, with a metamorphic peak at about 300 Ma. The rare monazite from the trondhjemitic sample yields evidence for crystallization age of about 275 Ma. In the interpretation of the in situ monazite ages obtained by the present study it is also necessary to take
HREE. The correlation between these elements is good for samples PB14 and GC30 but for sample GC13 only a poor correlation can be observed
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Fig. 7 Weighted-histogram representation of monazite age data for the paragneiss (a) and migmatite (b) samples. Each small bell-shaped curve represents the probability density function for one measurement. The dotted curve is the sum of all individual bellshaped curves. The thick curve represents the weighted mean age calculated by the statistical procedure
into account the lack of pre-Hercynian ages in all the analysed monazites, that would suggest an entirely Hercynian evolution of the rocks of the Aspromonte– Peloritani Unit. On the other hand, TIMS and ion-probe U–Pb zircon Archean to Neoproterozoic–Early Cambrian ages (Schenk 1990; Micheletti et al. 2007; Fiannacca et al. 2008, 2009) and U–Pb titanite and Ar–Ar hornblende Proterozoic ages (De Gregorio et al. 2003) are reported for magmatic and metamagmatic rocks of the same tectonic unit in both southern Calabria and north-eastern Sicily. By focusing first on the Hercynian evolution, the studied unmigmatized paragneiss and the migmatite yielded monazite ages of 298 ± 6 and 311 ± 4 Ma, respectively. The age of 298 ± 6 Ma provides for the Peloritani Mountains a scenario comparable to that already depicted
for the adjacent Aspromonte Massif of Southern Calabria, where Graeßner et al. (2000) dated the Hercynian metamorphic peak at 295 to 293 ± 4 Ma based on ID-TIMS U–Pb dating of monazite from similar amphibolite facies paragneisses. Combined geochronological and petrological data available for the whole southern Calabria–Peloritani Orogen suggests that low-P/high-T metamorphism overprinted the Barrovian one at about 300 Ma (Festa et al. 2004 and references therein). Graeßner and Schenk (1999) and Caggianelli and Prosser (2002) suggested that this lowpressure metamorphism was the result of combined decompression and T increase due to widespread granitoid plutonism. The above U–Pb ages were considered to represent the time of monazite growth and recrystallization driven by the heat input provided by the large
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Fig. 8 PbO versus ThO2* isochron diagram for the paragneiss (a) and migmatite (b) samples. The size of the X- and Y-bars of each data point represents the 2-sigma error, calculated from the counting statistics of the unknown and the standard. Data points marked with a square were not used for the calculation of the isochron and for the calculation of a weighted mean age
metaluminous to strongly peraluminous granitoid intrusions into the middle crust (Graeßner et al. 2000). This last Hercynian stage, occurring after the main deformational events, obliterated the previous ones almost everywhere and only rare relics allow constraining the prograde P–T path (e.g. Acquafredda et al. 2006; Angı` et al. 2010). Evidence for mineral and textural readjustment connected to thermal metamorphism and fluid–rock interaction is widespread in the rocks of the north-eastern Peloritani. In the studied samples post-tectonic crystallization of biotite and plagioclase, associated with local foam texture development or occurrence of feldspar replacement textures, represent examples of such processes. The age of 311 ± 4 Ma of the migmatite sample, about 10 Myr older than the age of the non-migmatitic paragneiss, might point to a somewhat earlier event, since it precedes the bulk of
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the Hercynian magmatism in the Calabria–Peloritani Orogen. This event might be framed in a similar or in a postemplacement genetic context of the adjacent trondhjemite body of Pizzo Bottino, which has a zircon SHRIMP magmatic crystallization age of 314 ± 3.5 Ma (Fiannacca et al. 2008). Both obtained ages, however, have to be referred to a thermal stage which postdated any blasto-deformational event, since the 314 ± 3.5-Myr-old trondhjemite cut discordantly the migmatitic foliation, providing therefore evidence for foliation development before 314 Ma. Therefore, monazite ages obtained for metasedimentary rocks from the Peloritanian segment of the Calabria– Peloritani Orogen are in agreement with the scheme proposed by Graeßner et al. (2000), explaining the metamorphic peak recorded in southern Calabria at about 300 Ma as the
Int J Earth Sci (Geol Rundsch) (2011) 100:107–123
result of combined decompression and thermal overprint related to magmatism. Based on available data, no clear evidence for an earlyHercynian event related to the thickening stage or to the early crustal melting processes leading to migmatite formation may be derived from the present monazite dating study. The data population of the apparent ages that range up to 340 Ma from meso-melanosomes of the studied migmatite (PB14) and up to 361 Ma in the non-migmatitic paragneiss (GC13) may only indicate monazite crystallization during prograde metamorphism. Trondhjemitic leucosomes produced by low-temperature H2O-present melting might be the best candidates to represent the earliest appearance of Hercynian melt in the Peloritani Mountains, and trondhjemitic leucosomes, interpreted as produced under prograde metamorphism, are reported from the lower crustal section of the Serre Massif (Fornelli et al. 2002). Nevertheless, leucosome monazite has not been dated because only one monazite grain was found in the studied sample. Different explanations can account for the monazite ages obtained for the studied migmatite, by considering that petrographic features and chemical data of these migmatites indicate a complex history involving both fluidsaturated and fluid-absent partial melting processes followed by retrogression and metasomatism (Fiannacca et al. 2005b). In particular, H2O-present melting was probably responsible for generations of quartz–plagioclase dominated leucosomes, whereas fluid-absent melting generated K-feldspar-rich granitic melt which locally appears to have mixed to variable extent with the trondhjemitic leucosomes. Thus, the monazite age of 311 ± 4 Ma could be related to the second event of partial melting experienced by the rocks, leading to the formation of ‘‘mixed’’ trondhjemite–granite leucosomes, as also reported for the Serre migmatites (Fornelli et al. 2002). Another possibility, supported by the widespread occurrence of feldspar replacement textures, might invoke monazite dissolution and recrystallisation directly caused by the infiltration of fluids released from the adjacent crystallizing magmatic Pizzo Bottino pluton. These interpretations are consistent with the depicted tectono-metamorphic evolution of the southern Calabria–Peloritani Orogen. Similar evidence may be obtained by comparison with adjacent segments of the southern Hercynian Belt, part of which was northeastern Sardinia before the post-Oligocene southeastern drifting of the Calabria and Sicily basements to their present position. In Sardinia, in situ Ar–Ar white mica ages and U–Pb zircon ages of 350–320 Ma (Di Vincenzo et al. 2004; Palmeri et al. 2004; Giacomini et al. 2006) have been obtained for the collision-related metamorphism. Metatexites with trondhjemitic leucosomes and Rb–Sr wholerock age of 344 ± 7 Ma (Ferrara et al. 1978) probably
119
represented the first occurrence of melt under prograde metamorphism in the basement of north-eastern Sardinia and, indeed, have been considered to reflect the collision stage or the beginning of exhumation (Giacomini et al. 2006 and references therein). Lower amphibolite facies mineral associations were produced during exhumation at 320–300 Ma, and finally, the last exhumation stages were marked by granitoid plutonism at about 310–290 Ma, involving high-temperature low-pressure metamorphic overprints. A similar tectono-metamorphic evolution, typical of continental collision chains, with crustal thickening followed by gravitative collapse, exhumation and granitoids emplacement is reported from many other segments of the European Hercynides such as Corsica (Menot and Orsini 1990), western and central Maures (Bellot et al. 2005); French Massif central (Roig and Faure 2000; Costa and Rey 1995; Rossi et al. 2006), and intra-Alpine massifs (Paquette et al. 1989; von Raumer et al. 1999; Rubatto et al. 2001). The absence of pre-Hercynian ages, suggesting a solely Hercynian single-stage metamorphic evolution, and of ages reliably constraining the Hercynian collisional peak in the studied samples might be ascribed to monazite recrystallization. Similar situations are reported from many highgrade rocks (e.g. Montel et al. 2000; Berger et al. 2005; Braun and Appel 2006; Braun et al. 2007) where monazite crystals totally included in garnet or quartz preserved older ages compared to monazite crystals from the rock matrix or from cleavage-bearing, fluid-accessible, mineral phases such as orthopyroxene, kyanite or biotite. Communication of monazite with other relevant Pb-containing phases in metapelite (e.g. apatite, plagioclase) has been invoked as a cause for monazite signature age resetting, independently from further textural location or difference in grain size of the monazite grains (Berger et al. 2005). Dissolution/precipitation in the presence of a fluid phase is also commonly considered a possible mechanism for partial to complete resetting of the monazite isotope system (Seydoux-Guillaume et al. 2002) and complete dissolution and reprecipitation of monazite indeed has been frequently ascribed to heat pulse and/or alteration by fluids released by crystallizing magmatic bodies (e.g. Zeh et al. 2003 and references therein). In the studied metasedimentary rocks monazite occurs as an inclusion in biotite or as a matrix phase, both types of textural locations making monazite vulnerable to both recrystallisation and fluid-induced dissolution/reprecipitation leading to the resetting of the U–Pb system. Moreover, although most of the youngest individual ages occur in the outermost part of the grains, there is no systematic relation between age data and the position of the analytical spots within the monazite grains. Finally, diffuse alteration
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affecting monazite along cracks may provide further support for monazite resetting late in the metamorphic history of the rocks. The monazite study of the trondhjemite sample did not allow to constrain the emplacement age of the trondhjemite pluton and to obtain clear evidence for the occurrences of different crystallization stages of magmatic and metasomatic monazite, due to the very low amount and small size of monazite in the sample analysed. Nevertheless, these results provide preliminary information about the age of monazite growth in the Peloritanian trondhjemitic rocks and in conjunction with data provided by recent studies (Fiannacca et al. 2005a, 2008), they add new facts useful to gain understanding of the nature and origin of these rocks. First, the composition of trondhjemite monazite plots in a very restricted field of the ternary diagram 2REEPO4–CaTh(PO4)2–2ThSiO4 (Fig. 3a), indicating a cheralite-enriched composition with low huttonite component. Similar monazite, with cheralitic component greater than that normally found in granites (6–18%, according to Bea 1996), has been mainly reported from, but is not restricted to, strongly peraluminous S-type granites (Fo¨rster 1998). Fo¨rster (1998) also reports that cheralite appears to crystallize during late-stage processes from fluid-rich residual liquids rather than during early magmatic crystallization. Cheralite-rich monazite is also found as a secondary phase, for example associated to chloritization affecting peraluminous granites (Poitrasson et al. 1996). Second, the preliminary chemical ages obtained from dating of sample GC30 seem to indicate that monazite crystallization occurred at about 275 Ma and, in any case, not earlier than 290 Ma. This data strongly conflicts with recent U–Pb SHRIMP dating, which reliably documents crystallization of magmatic zircon at 314 ± 3.5 Ma in a sample from the same trondhjemite body (Fiannacca et al. 2008). At present, the only possible explanation appears that, if the chemical ages really reflect crystallization of monazite at about 275 Ma, this crystallization should have occurred under post-magmatic conditions, possibly during the same metasomatic event invoked to account for the unusual petrographical and geochemical features of the trondhjemites (Fiannacca et al. 2005a). This interpretation appears to be consistent with the cheralite-rich composition of studied monazite.
Conclusion The present study reports the first ‘‘in situ’’ chemical ages of monazite from two metasedimentary rocks and one
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trondhjemite from a former middle crustal sector of the Calabria–Peloritani Orogen. The data obtained for one biotitic paragneiss and one migmatite shows that they shared the same Hercynian metamorphic evolution, climaxing at ca. 300 Ma, in a low-pressure–high-temperature event, which occurred in a context of combined decompression and magmatism, a scenario which is generally accepted for the entire Calabria–Peloritani Orogen at that time. The age of 311 ± 4 Ma of the migmatite sample, about 10 Ma older than that of the non-migmatitic paragneiss, might point to a somewhat earlier event. This might be framed in the same genetic, or post-emplacement, context of the adjacent trondhjemite body of Pizzo Bottino, having a zircon SHRIMP magmatic crystallization age of 314 ± 3.5 Ma (Fiannacca et al. 2008). No clear evidence has been found for both metasedimentary samples, which allows temporal constraints on the previous Barrovian metamorphic peak linked to the thickening stage. Some individual dates, however, are on the high age side of the Gauss distribution and are either without any geological significance or represent relics of monazite crystallization during prograde metamorphism. In either case, the absence of monazite ages older than 361 Ma indicates a solely Hercynian single-stage metamorphic evolution for the metasedimentary rocks of the northern Peloritani Mountains. Nevertheless, monazite occurrence in textural locations favourable to isotopic resetting, namely as a matrix phase or as an inclusion in cleavage-bearing mineral phases, leaves open the possibility that the record of older events could have been completely erased during the Hercynian metamorphic evolution of the rocks. The results of in situ monazite chemical dating of the trondhjemite sample have to be considered preliminary at present, due to the very low amount and small size of analysed monazites. There results represent, however, the first chemical and geochronological data of monazites from trondhjemites of the Peloritani Mountains and, they therefore, give a contribution to the knowledge of these rocks. The chemical ages obtained in this work indicate the starting of monazite crystallization after 290 Ma, with a weighted mean age of 275 ± 4 Ma. Although this data has to be considered as provisional, it would indicate a crystallization age of the monazite too far ahead in time compared to the real age of the magmatic crystallization, at ca. 314 Ma, of the zircon from the same plutonic body, to be considered a magmatic age. Consequently, monazite ages obtained for the studied sample are better interpretable as related to a post-magmatic stage of subsolidus crystallization, in accordance with the metasomatic model proposed by Fiannacca et al. (2005a) to explain the origin of these peculiar trondhjemites.
Int J Earth Sci (Geol Rundsch) (2011) 100:107–123 Acknowledgments We thank A. Berger, an unknown reviewer and the topic editor I. Braun for critical comments and suggestions that helped significantly to improve the manuscript.
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