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Journal of Petrology Volume 42 Number 2 Pages 355-375 2001
© Oxford University Press 2001
Ilmenite as a Source for Zirconium during High-grade Metamorphism? Textural Evidence from the Caledonides of Western Norway and Implications for Zircon Geochronology
1GEOLOGICAL SURVEY OF NORWAY, LEIV EIRIKSSONS VEI 39, N-7491 TRONDHEIM, NORWAY
2MINERALOGISK GEOLOGISK MUSEUM, UNIVERSITY OF OSLO, N-0562 OSLO, NORWAY
3SWEDISH MUSEUM OF NATURAL HISTORY, SE-104 05 STOCKHOLM, SWEDEN
Received August 16, 1999; Revised typescript accepted June 5, 2000
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONMETAMORPHISM IN THE LINDAS...METHODSGRANULITES AND AMPHIBOLITESSIMS Th-U-Pb data on...ECLOGITESIMS Th-U-Pb data on...DISCUSSIONCONCLUSIONSREFERENCES |
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The Proterozoic Lindås Nappe, part of the Caledonides of western Norway, was affected by penetrative Sveconorwegian granulite-facies metamorphism, followed by a fluid-driven eclogite- and amphibolite-facies Caledonian overprint, spatially restricted along fractures and shear zones. In mafic granulites and amphibolites, a luminescent anhedral zircon overgrowth, which gives an average age of 924 ± 58 Ma (Th/U = 0·52; secondary ion mass spectrometry data), surrounds a magmatic zoned core with an age of 952 ± 32 Ma (Th/U = 1·27). In the granulites, a continuous rim of zircon or a discontinuous corona of
10 µm rounded to flat zircon crystals is observed at the outer margin of ilmenite grains. Baddeleyite and srilankite (Ti2ZrO6) blebs are reported around ilmenite included in feldspar or pyroxene. Baddeleyite is interpreted as an exsolution product from magmatic ilmenite, whereas srilankite, the zircon corona around ilmenite and the luminescent zircon overgrowth were formed as reaction products during granulite-facies metamorphism. Textures suggest that magmatic ilmenite was a main source of Zr to form metamorphic zircon. In massive amphibolites, relic ilmenite grains are surrounded by a corona of titanite and a discontinuous corona of micro-zircons. Amphibolite-facies overprint is not associated with any significant growth or dissolution of zircon. An unsheared eclogite displays a zircon population with a euhedral oscillatory zoned overgrowth giving an age of 455 ± 29 Ma (Th/U 
0·13). A corona of micro-zircon grains is observed at some distance around rutile, and locally these zircons show a prismatic overgrowth. A specific low-Th zircon growth event is related to eclogite-facies forming reactions, involving breakdown of a two-pyroxene + garnet + plagioclase + ilmenite assemblage to form a garnet + omphacite + rutile assemblage in the presence of a fluid. The low Th content of this zircon probably stems from the coeval precipitation of clinozoisite. This oscillatory zoned zircon records fluid infiltration and coeval eclogitization in the crust. KEY WORDS: eclogite; granulite facies; ilmenite; zircon; Caledonides
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INTRODUCTION |
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In high-grade metamorphic rocks, zircon generally has complex internal structures (Black et al., 1986
; Corfu et al., 1998; Whitehouse et al., 1999), reflecting the resistant but complex behaviour of this mineral during metamorphism [see reviews by Heaman & Parrish (1991) and Mezger & Krogstad (1997)]. The use of high-precision or high-spatial-resolution U–Pb analytical methods provides reliable age information for the internal domains of zircon crystals (Krogh, 1993; Bowring & Williams, 1999). Dating metamorphism thus involves a decision as to which domains of the zircon formed during particular episodes of metamorphism. Typically this decision relies on characterization of the morphology and internal structure of zircon populations, and this is generally the weak step in the interpretation of the data. Rigorous and detailed interpretation of geochronological data requires that we relate the formation of zircon domains to well-defined metamorphic changes or reactions in the rock (Fraser et al., 1997). A number of mechanisms have been proposed for the formation of zircon during metamorphism:
- annealing or lattice reorganization occurs in inherited metamict zircon during heating above 600°C and is associated with episodic Pb loss [summary given by Mezger & Krogstad (1997)].
- Secondary crystallization corresponds to the replacement of an inherited zircon lattice by a lattice of distinct composition, possibly by reaction with a metamorphic fluid. The process can follow pre-existing structures of the zircon or takes the form of reaction fronts migrating from the surface of the grain (Pidgeon et al., 1998; Schaltegger et al., 1999; Vavra et al., 1999).
- Precipitation of metamorphic zircon can result from solid-state reactions involving breakdown of magmatic or metamorphic Zr-bearing phases (Davidson & van Breemen, 1988; Fraser et al., 1997). Breakdown of any modally abundant mineral containing tens of ppm of Zr allows for significant zircon formation.
- Precipitation of zircon is possible in a partial intergranular melt, itself enriched in Zr by partial resorption of inherited zircon or other Zr-bearing phases (Roberts & Finger, 1997; Vavra et al., 1999). Garnet, amphibole, pyroxene and ilmenite can represent significant repositories for Zr in magmatic and metamorphic rocks, which may be made available during metamorphism and lead to zircon precipitation (McLelland & Chiarenzelli, 1990; Fraser et al., 1997; Scoates & Chamberlain, 1997).
Eclogite-facies rocks usually contain rounded to multifaceted zircons commonly with a voluminous inherited core surrounded by an overgrowth (Peucat et al., 1982). This overgrowth possibly formed during eclogite-facies metamorphism. The occurrence of inclusions of high-pressure index minerals in zircon such as rutile, omphacite, coesite or diamond is good evidence for zircon growth during the eclogite-facies event itself (Krogh et al., 1974; Claoué-Long et al., 1991; Ames et al., 1993; Rowley et al., 1997). In the absence of such evidence, eclogite-facies zircon is considered as poorly zoned to sector zoned, and characterized by a distinctly low Th/U ratio (Claoué-Long et al., 1991; Creaser et al., 1997; Rowley et al., 1997; Brueckner et al., 1998; Hacker et al., 1998). Oscillatory zoned zircon, otherwise regarded as typically magmatic, has been rarely assigned to eclogite-facies metamorphism in the presence of hydrous fluid (Rubatto et al., 1998).
In the Lindås Nappe, Bergen Arc, Caledonides of western Norway (Fig. 1), a Proterozoic plutonic complex is affected by Sveconorwegian–Grenvillian granulite-facies metamorphism and by Caledonian eclogite- and amphibolite-facies overprint. The Caledonian overprint is spatially restricted along fractures and shear zones. Field relationships clearly indicate that amphibolitization or eclogitization of the anhydrous protolith was triggered by fracturing and fluid infiltration. The reactions were arrested where fluid was not available during metamorphism, leaving large volumes of unaffected metastable granulite (Austrheim, 1987; Boundy et al., 1992; Austrheim et al., 1997). Such relationships are found in many high-pressure terrains (Austrheim, 1998). As metamorphic transitions are observed in situ, this occurrence provides an opportunity to address fundamental questions related to metamorphism. In this paper, we focus on the petrology of Zr minerals, namely baddeleyite, srilankite and zircon, in mafic lithologies and their relationships with ilmenite and rutile, using scanning electron microscopy (SEM) and secondary ion mass spectrometry (SIMS). The objective is to relate metamorphic changes in Ti minerals and Zr minerals, to characterize the formation of metamorphic zircon.
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METAMORPHISM IN THE LINDÅS NAPPE |
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The Bergen Arc comprises a set of arcuate nappes, including the Lindås Nappe (Fig. 1) (Kolderup & Kolderup, 1940
). The Lindås Nappe mainly consists of a Proterozoic anorthosite–mangerite–charnockite–granite (AMCG) suite and banded gneiss complexes, affected by Sveconorwegian granulite-facies metamorphism. Granulite-facies assemblages provide P–T estimates of <1 GPa and 800–850°C (Austrheim & Griffin, 1985). The level of intrusion of the AMCG suite and the P–T path it followed from the magmatic crystallization to the granulite-facies overprint is difficult to assess. Nevertheless, the occurrence of a penetrative granulite-facies deformation, of widespread coarse-grained garnet or scapolite-bearing mineral assemblages (Moecher & Essene, 1991), and of common corona textures around relic magmatic minerals (mainly olivine) (Griffin, 1972; Cohen et al., 1988) support the inference that the original magmatic assemblages were thoroughly overprinted and almost completely replaced by granulite-facies ones. The granulite-facies event was followed by regional cooling at 920–900 Ma (Rb–Sr and Sm–Nd mineral isochron ages ranging from 923 ± 7 Ma to 896 ± 31 Ma) (Cohen et al., 1988; Burton et al., 1995). Phlogopite Rb–Sr ages of 882 ± 9 to 835 ± 7 Ma attest to the local retention of Proterozoic ages in this mineral (Kühn et al., 2000).
The Lindås nappe shows partial eclogite- and amphibolite-facies Caledonian overprint. Eclogites are most abundant on the island of Holsnøy (Fig. 1). There, three stages of progressive and arrested eclogitization of the granulite protolith are recognized (Austrheim & Griffin, 1985; Austrheim, 1987; Jamtveit et al., 1990; Boundy et al., 1992):
- eclogitization was initiated along brittle fractures oblique to the granulite foliation; it progressed as a reaction front migrating away from the fractures. The development of euhedral eclogite-facies hydrous minerals (namely, phengite, clinozoisite and amphibole), within and along the fractures, attests to the introduction of fluid. Fluid probably played an important role in the fracturing of the protolith and development of the eclogite-facies assemblages.
- Discontinuous, thin (<2 m thick), ductile shear zones with eclogite-facies mineralogy developed along the fractures.
- Minor shear zones merged to form broad shear zones containing rotated granulite-facies blocks.
The largest shear zones are up to 100 m thick and several kilometres long and contain few granulite relics. Eclogite-facies assemblages record P–T conditions of 1·8–2·1 GPa and 700°C (Jamtveit et al., 1990). The development of the amphibolite-facies overprint occurred in a similar fashion (Austrheim & Robins, 1981). It started with static conversion of the mineral assemblages and developed into amphibolite-facies shear zones. Amphibolite-facies assemblages commonly replace eclogite-facies assemblages, implying that an amphibolite-facies overprint followed the eclogite-facies one. Although eclogitization and amphibolitization are associated with fluid introduction and changes in petrophysical properties, the metamorphic transitions can be considered isochemical (Rockow et al., 1997).
The timing of eclogite-facies metamorphism is estimated by titanite and epidote 238U–206Pb ages of 465–445 Ma, from eclogitic marbles (Boundy et al., 1997). Hornblende 40Ar/39Ar plateau ages of 455 ± 2 and 448 ± 4 Ma, and muscovite 40Ar/39Ar plateau ages of 433 ± 1 to 429 ± 1 Ma have been interpreted as cooling ages after the eclogite- and amphibolite-facies events (Boundy et al., 1996).
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METHODS |
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The petrography of 30 samples was investigated in thin section with both optical microscopy and SEM, and the results from 12 samples (Table 1) are reported in this paper. Phases were identified with the help of backscattered electron imaging (BSE) and semi-quantitative energy dispersive X-ray spectrometry (EDXS). BSE and cathodoluminescence (CL) images were taken from selected areas in thin section to study zircon habits and petrographical relationships in situ. Quantitative microanalyses of srilankite were acquired with a Cameca Camebax microprobe (University of Oslo) at the operating conditions of 15 kV accelerating voltage and 20 nA beam current. Zircon with 2 wt % HfO was used as standard for Hf, Zr and Si. Vanadite was used as standard for V and the VKß peak was counted to avoid overlap with Ti. Data reduction was carried out using the Cameca PAP software package.
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Zircon was extracted from five of the samples using a Wilfley water table, heavy liquids and a magnetic separator. Zircon was analysed for Th, U and Pb by SIMS using the Cameca IMS 1270 instrument at the Nordsim laboratory, Swedish Museum of Natural History, Stockholm. Analyses were conducted following the method outlined by Whitehouse et al. (1997, 1999). Two types of analyses were carried out on epoxy grain mounts: analyses of cross-sections polished to approximately half-thickness, and surface analyses. For surface analyses, the epoxy mount was polished for a short time (30 s with 1 µm diamond spray) so that the surface of the zircon grains was exhibited at the surface of the mount and slightly polished. Analyses of cross-sections (Table 2) were performed after systematic CL imaging of 30–70 crystals per sample, to select the most suitable analysis locations. Analyses were conducted on two samples (BH10, BH11) with an 4 nA O2- beam and a spot size of 40 µm, and on a third sample (BH2) with an 2 nA beam and a spot size of 20 µm. The Kipawa reference zircon was used for calibration with an age of 993 Ma (Stern, 1997). Surface analyses (Table 3) were carried out with an 4 nA beam and a spot size of 40 µm, using the Geostandard 91500 reference zircon with an age of 1065 Ma (Wiedenbeck et al., 1995). Because of the low U and Pb concentrations in a large proportion of the analysed zircon grains, the 235U–207Pb and 207Pb/206Pb ages are critically dependent on 207Pb counting statistics and the handling of the common Pb signal (Pbc). Only the 238U–206Pb age was thus selected for discussion. A Pbc correction was applied on the basis of the 204Pb signal (e.g. Zeck & Whitehouse, 1999) with present-day isotopic composition following Stacey & Kramers (1975). Weighted average ages were calculated at the 95% confidence level using a subroutine of the Isoplot program (Ludwig, 1999). Application of the standard weighted average procedure results in a mean square weighted deviation (MSWD) higher than two for most groups of analyses, indicating scatter in excess of analytical error. As a result, an algorithm that partitions the error of the analyses into their analytical component and an excess component of constant external error has been used to calculate average values.
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Zircon U–Pb geochronology was conducted on the same sample set by isotope dilution thermal ionization mass spectrometry (ID-TIMS). These data are reported in a companion paper focusing on the geochronology and tectonic evolution of the Lindås nappe (Bingen et al., 2001), and are quoted in the following discussion. ID-TIMS analyses demonstrate that homogeneous zircon fractions in these samples are concordant and have a very low Pbc content. This gives us confidence in the accuracy of the SIMS 238U–206Pb ages used below.
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GRANULITES AND AMPHIBOLITES |
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Mineral assemblages
Nine samples of granulite were selected from the jotunite to mangerite rock units on the islands of Radøy and Holsnøy (Fig. 1; Table 1). Typically, these units have SiO2 of 46–53 wt % (Austrheim & Mørk, 1988
; Rockow et al., 1997). The samples display millimetre-scale, equigranular, poorly to moderately oriented, granoblastic textures. The sampling includes granulites (BH3.9, M262), hornblende granulites (BH10, BH1B3), garnet granulites (BH5, HA61/89a, HA55/74), a scapolite granulite (BH1B2) and a quartz granulite (HA83/73). All samples contain two pyroxenes, a perthitic to mesoperthitic plagioclase, ilmenite, apatite and zircon. Part of the samples contains sulphides and magnetite. Ilmenite is rounded to anhedral and consists in coarse exsolution lamellae of haemo-ilmenite and ilmeno-haematite (Fig. 2). Apatite typically includes small needle-shaped inclusions of monazite (1–10 µm wide) (Austrheim et al., in preparation).
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Although efforts have been made to collect the best preserved granulite samples, almost all localities in the Lindås nappe are affected by at least some Caledonian overprint. For example, sample BH10 comes from the centre of an
5 m thick lens of granulite surrounded by amphibolite. Centimetre-wide fractures associated with the development of green amphibole + biotite + epidote occur in the lens and the sample. A garnet + omphacitic pyroxene (5·0 wt % Na2O) association is locally observed along the fractures, suggesting a first phase of high-pressure overprinting. At microscopic scale, a zone of fine-grained secondary minerals and symplectites of 10–60 µm thickness occurs at many grain boundaries in sample BH10 and all other granulite samples [illustrations have been given by Austrheim & Robins (1981) and Austrheim et al. (1997)]. These minerals grew at the expense of feldspar and pyroxenes, and are either eclogite- or amphibolite-facies minerals. Both types of alteration can occur in the same sample. The eclogite-facies alteration is characterized by an omphacite + amphibole symplectite, mainly located at the margin of pyroxene aggregates, and by a garnet + quartz symplectite, mainly located around ilmenite and garnet grains. Kyanite needles are common in feldspar. The amphibolite-facies alteration comprises an amphibole ± quartz ± plagioclase symplectite at the margin of pyroxene grains, and fine-grained biotite or a biotite + quartz symplectite at the margin of ilmenite grains. The amphibole- and eclogite-facies alteration symplectites contain minor amounts of fine-grained (<10 µm) opaque minerals. These give a bright signal on BSE images (Fig. 2a and g), and are mostly Fe-rich oxide and sulphide minerals, presumably haematite and pyrite. Apatite grains commonly display a clear external margin devoid of monazite inclusions, and are surrounded by a rim of fine-grained allanite (10 µm thick).
Two samples of amphibolite were selected on Radøy. Sample BH11 was collected from the same outcrop as the granulite BH10, and comes from a shear zone of 3 m thickness. It is totally converted to an amphibolite gneiss showing a two-feldspar + amphibole + biotite + epidote + titanite assemblage. The massive amphibolite BH8 was collected in a lens of jotunite of 10 m thickness, showing static amphibolite-facies overprint. The sample contains some garnet, and displays relics of coarse-grained granulite-facies perthitic plagioclase and ilmenite. Perthitic plagioclase relics are situated inside aggregates of two feldspars. Radially oriented biotite and amphibole grains commonly surround pyroxene pseudomorphs and ilmenite grains. Ilmenite is typically surrounded by a corona of polycrystalline titanite of 100 µm width, clearly growing at the expense of ilmenite, and an outer corona of biotite.
Textures involving zircon
Zircon has identical habit in the granulites and amphibolites. It is the main Zr mineral, and is randomly distributed in the rock as rounded to anhedral grains ranging from 40 to 500 µm in size. Prismatic grains are less common, and have been observed included in feldspar. In the granulites, coarse-grained zircon grains (>60 µm) display a specific habit when they are located in contact with ilmenite or more rarely with apatite. There, they have the shape of a ‘hat’, with the flat base of the ‘hat’ forming the contact with ilmenite or apatite (Fig. 3).
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A widespread feature of the granulites is the occurrence of small, rounded to flat zircon crystals, typically of 10 µm thickness, at the external edge of ilmenite grains. These zircon crystals commonly form a discontinuous corona, identified by a positive BSE contrast (Fig. 2a). Locally, a thin continuous cover of zircon, up to 500 µm in length, is observed on ilmenite (Fig. 2b and d). Locally, a discontinuous corona of small zircon grains or a continuous cover of zircon grades into the flat base of a coarse (>60 µm) ‘hat’-shaped zircon grain in contact with ilmenite (Figs 2c and 3). A corona of micro-zircons is not recorded at the surface of all ilmenite grains, but the spatial association between micro-zircon grains and ilmenite is well defined in the granulites. It is not feasible to demonstrate that all micro-zircon grains are associated with ilmenite, and so a minor proportion of them may also be scattered randomly in the remaining volume of the rock.
In the amphibolite sample BH8, for which amphibolite-facies overprint is static, a discontinuous corona of small (10 µm) rounded to flat zircon grains is spatially associated with relic ilmenite. The zircon corona occurs at the contact between the titanite corona growing at the expense of ilmenite, and the outer biotite corona (Fig. 2e).
Baddeleyite and srilankite
Srilankite (Ti2ZrO6) crystals are observed at the margin of ilmenite in five granulite samples. They form rounded to prismatic grains, up to 50 µm in size, situated at the edge of ilmenite, typically included by their host (Fig. 2f and g). They were observed in the granulites BH1B2 and BH3.9, the garnet granulites 55/74 and HA61/89, and the scapolite granulite BH1B2. Microprobe analyses in two samples (Table 4) yield a Zr/Ti ratio close to a stoichiometric value of one-half, which is characteristic of srilankite as described by Willgallis et al. (1983). Srilankite occurrence is limited to the ilmenite grains surrounded by mesoperthitic feldspar and to locations showing limited post-granulite alteration at grain boundaries. Gahnite-rich (ZnAl2O4) spinel inclusions commonly occur in the same grains. Srilankite crystals show no systematic spatial relationship with exsolution lamellae of haemo-ilmenite and ilmeno-haematite.
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Baddeleyite (ZrO2) crystals of 5–30 µm in diameter are detected at the margin of ilmenite in five granulite samples. In the granulites BH1B2 and BH3.9, and the scapolite granulite BH1B2, baddeleyite is associated with srilankite at the margin of ilmenite grains included in feldspar. Baddeleyite occurs in such a way that it is not observed in direct contact with ilmenite (Fig. 2h and i): baddeleyite occurs at the external margin of the ilmenite grain but is separated from it by a rim of srilankite, or baddeleyite is included in srilankite.
Two rare occurrences of baddeleyite are reported in which baddeleyite is in direct contact with ilmenite. In the granulite HA83/73, a flat baddeleyite crystal (40 µm long) is observed at the surface of an ilmenite grain included in clinopyroxene (Fig. 2j). In the granulite M262, a 30 µm rounded baddeleyite inclusion in an ilmenite grain is associated with a thin continuous cover of zircon at the surface of this ilmenite grain (Fig. 2k). The thickness of the zircon cover seems to decrease away from the baddeleyite crystal.
CL images of zircon
CL images were taken of polished sections of zircon extracted from two granulites and two amphibolites. The zircon populations of the granulite BH10 and nearby amphibolite BH11 are identical. Crystals are anhedral to prismatic, and commonly larger than 200 µm width. A large majority of zircon crystals contain a main core characterized by euhedral oscillatory to sector zoning and an anhedral homogeneous and luminescent overgrowth of variable thickness (Fig. 4a–c). The core–overgrowth interface commonly displays embayments that truncate the internal structure in the core. A non-luminescent (dark) oscillatory-zoned inner core is sometimes observed (Fig. 4d). The inner-core–core interface is rounded. ID-TIMS analyses of deeply abraded fractions of BH10 and BH11 give equivalent ages of 951 ± 2 and 951 +10/–4 Ma. These ages reflect the formation of the oscillatory zoned core (Bingen et al., 2001).
In the garnet granulite BH5, zircon is rounded to anhedral. It is characterized by a small non-luminescent core surrounded by a very thick, homogeneous anhedral luminescent overgrowth, which generally represents >90% of the volume of the zircon (Fig. 4e). ID-TIMS analyses of six fractions give overlapping concordant ages at 929 ± 1 Ma, interpreted as the age of formation of the dominant luminescent overgrowth (Bingen et al., 2001).
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Sample BH8 contains abundant zircons of rounded to thick prismatic shape (commonly of >200 µm width). The zircons are characterized by a very luminescent core (low U content), with poorly defined oscillatory zoning, and an anhedral luminescent and homogeneous overgrowth (Fig. 4f). ID-TIMS analyses of two fractions of deeply abraded coarse-grained zircons show very low U content (6–17 ppm) and give an age of 957 ± 11 Ma. These analyses mainly represent the formation of the very luminescent core. Two fractions of small rounded zircons have a higher U content (29–42 ppm) and are concordant at 933 ± 2 Ma. These fractions are probably dominated by the moderately luminescent overgrowth, and their age reflects the crystallization of this overgrowth (Bingen et al., 2001
). ‘Hat’-shaped zircons are common in the mineral separates of the four samples. In the granulites BH10 and BH5 these were probably in contact with ilmenite in the rock, and in the amphibolites BH11 and BH8, they are probably inherited from the granulite stage, during which they were in contact with ilmenite. Special care has been taken to mount and take CL pictures of ‘hat’-shaped zircons (Fig. 4b, e and f). These zircons show a rounded core having similar internal structure and contrast to the core in other grains of the sample (except for sample BH5 where cores are rare and small), and the ‘hat’-shaped overgrowth is homogeneous and luminescent. The orientation of the flat base of the ‘hat’ does not correspond to any of the main crystallographic orientations of the core.
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SIMS Th–U–Pb data on zircon |
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Cross-sections
The combination of CL imaging technique and conventional ID-TIMS zircon dating in different samples allows resolution, without ambiguity, of the formation of the oscillatory to sector zoned core at 951 ± 2 Ma, and the formation of the luminescent overgrowth at 933 ± 2 to 929 ± 1 Ma. SIMS data were acquired in the granulite BH10 and nearby sheared amphibolite BH11 to further characterize the zircon population. SIMS analyses of the zircon yield similar results for the two samples and can be described together (Figs 5 and 6). The non-luminescent inner core is rich in U (337–832 ppm, average 551 ppm), and has an average Th/U ratio of 0·54 (0·36–0·68). Five analyses give a weighted average 238U/206Pb age of 950 ± 26 Ma (MSWD = 2·5). The main zoned core has an average U content of 88 ppm (54–113 ppm), and a Th/U ratio of 1·27 (1·03–1·48). Sixteen analyses give an identical age of 952 ± 32 Ma (MSWD = 9·2). The luminescent anhedral overgrowth is poor in U (12–30 ppm; average 23 ppm) and has a Th/U ratio of 0·53 (0·40–0·74). Nine analyses yield an age of 924 ± 58 Ma (MSWD = 3·5). No Caledonian overgrowth can be detected in samples BH10 and BH11.
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SIMS analyses were carried out in ‘hat’-shaped zircon grains in sample BH10. The zoned rounded core gives a Th/U ratio of 1·3 (n = 3,
970 Ma) similar to that of the oscillatory to sector zoned core in other zircons in this sample. The ‘hat’-shaped luminescent overgrowth yields a Th/U ratio of 0·44 (n = 4, 960 Ma) similar to that of the luminescent anhedral overgrowth around other zircons in the sample (Table 2).
Surface data
SIMS analyses of the surface of zircons were performed in the granulites BH5 and BH10 and in the amphibolites BH8 and BH11, to estimate the age of the youngest overgrowth. Surface analyses yield variable U contents of 57 ppm for BH10 (n = 1), 23–32 ppm for BH5 (n = 5), 7–22 ppm for BH11 (n = 6), and 6–8 ppm for BH8 (n = 5). The Th/U ratio is very homogeneous around an average value of 0·23 (0·15–0·29, n = 17). The age ranges from 1109 ± 120 to 804 ± 74 Ma. Measurements are imprecise owing to low U contents (Figs 5 and 6). They are probably affected by Pb or U loss and possibly by systematic calibration errors, especially for low U analyses. Nevertheless, the lack of a Caledonian age signature in the four samples can hardly be fortuitous and points to the absence of a thin Caledonian overgrowth at the surface of zircon grains in the granulites and amphibolites.
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ECLOGITE |
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Mineral assemblage
Sample BH2 is a massive eclogite collected from Holsnøy, close to the granulite BH5 (Fig. 1, Table 1). The sample does not show an oriented structure so it is interpreted as the result of static eclogitization of the surrounding granulite. The sample displays a coarse-grained assemblage of garnet + omphacite (6·5 wt % Na2O), with minor blue–green amphibole (4·5 wt % Na2O), phengite, epidote, apatite, rutile, zircon and quartz. Relics of granulite-facies garnet (7·6 wt % MgO) occur in the core of eclogite-facies garnet (3·9 wt % MgO) [similar to textures described by Erambert & Austrheim (1993)
]. No relic ilmenite remains and only eclogite-facies rutile is observed. The sample does not show apparent evidence for post-eclogite facies alteration or decompression melting.
Zircon
In the eclogite BH2, zircon occurs as prismatic and rounded crystals. Zircon grains are smaller and less abundant than in the nearby granulite BH5. A discontinuous corona of small and flat zircons is commonly observed around rutile; it occurs at 20–100 µm from the rutile crystal and is included in eclogite-facies garnet (Figs 7 and 8).
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In the mineral separate, zircon occurs as euhedral well-terminated prisms, commonly multifaceted, as flat grains, and as anhedral grains. Polycrystalline aggregates of small crystals (<60 µm) are widespread (Fig. 9). CL imaging shows a variably luminescent core (Fig. 4g) with a euhedral overgrowth (Fig. 4g–m). The overgrowth is variably luminescent and commonly oscillatory zoned (Fig. 4h–m). The grains showing the best euhedral habit commonly display a small core (<10 µm) with a rounded to flat shape (Fig. 4j–m). In the polycrystalline aggregates, several small cores forming a row are commonly observed (Fig. 4h and l).
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ID-TIMS analyses of five fractions of zircon with a variously developed prismatic overgrowth give a discordia line with intercepts at 930 and 460 Ma. Regression of these fractions with the six fractions of the nearby granulite BH5 yields a lower intercept age of 456 ± 7 Ma. This age is interpreted as the age of formation of the euhedral overgrowth (Bingen et al., 2001).
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SIMS Th–U–Pb data on zircon |
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Cross-sections
In the eclogite BH2, SIMS analyses (Figs 5 and 6) of the core of large zircons (>60 µm) give a variable U content, ranging from 80 to 592 ppm, and a variable Th/U ratio (0·31–0·89). The age ranges from 777 ± 32 to 934 ± 36 Ma. The variability of morphology and chemistry of the core indicates that it represents a mixed population of magmatic and/or metamorphic origin, which is coeval with or younger than the zircons in the nearby granulite-facies rocks. Some analyses are possibly affected by Pb loss.
Concentrations of U, Th and Pb in the oscillatory zoned euhedral overgrowth are commonly below the detection limit of the SIMS method, especially in the poorly luminescent areas. Therefore, only two spot analyses could be acquired in this overgrowth where it is poorly luminescent, and nine analyses were acquired where it is more luminescent. All analyses considered together, the euhedral overgrowth has a U content between 0·4 and 152 ppm, and a low Th/U ratio below 0·13. It gives an age ranging from 521 ± 48 to 310 ± 72 Ma (Figs 5 and 6) with an average value of 428 ± 42 Ma. The three data points below 400 Ma are probably unreliable as a result of low U content and possible Pb loss. Discarding these points, eight analyses of the euhedral overgrowth give an average age of 455 ± 29 Ma (MSWD = 4·7), equivalent to the ID-TIMS age of 456 ± 7 Ma.
Surface data
SIMS analyses of the surface of zircons in the eclogite sample give a U content of <25 ppm and a Th/U ratio of <0·05. The calculated age is imprecise and ranges from 595 ± 236 to 242 ± 30 Ma (n = 6). The data indicate that a Caledonian overgrowth is ubiquitous in this sample. The surface of the overgrowth is variably affected by post-Caledonian Pb loss.
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DISCUSSION |
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Srilankite formation during granulite-facies metamorphism
The crystallographic lattice of ilmenite can accommodate Zr in significant amounts. Oxidized ilmenite containing up to 1 wt % ZrO2 is reported in the groundmass of kimberlitic magmas (Tompkins & Haggerty, 1985
). In magmatic systems, KD (Zr)ilm/liq ranges from 0·3 for basaltic magmas (McCallum & Charette, 1978) to 0·5–10 for rhyolitic magmas (Ewart & Griffin, 1994). Zr is thus a compatible element in ilmenite during magmatic crystallization, and magmatic ilmenite possibly represents a significant reservoir of Zr in mafic lithologies, together with amphibole, if present.
Baddeleyite is a widespread magmatic mineral in mafic rocks. In dykes and sills, it commonly occurs both as discrete crystals formed from interstitial, late-stage melts and as tiny blebs (<10 µm) at the margin of ilmenite (Heaman & LeCheminant, 1993). Lamellae of baddeleyite are reported inside ilmenite from a tholeiitic sill (Naslund, 1987). There, they are interpreted as a subsolidus feature, implying that the Zr content can be high in magmatic ilmenite and that baddeleyite can exsolve from the ilmenite lattice during subsolidus cooling of the rock. During granulite-facies metamorphism, baddeleyite is generally converted to zircon, as a response to increased silica activity during metamorphism (Davidson & van Breemen, 1988).
Srilankite was originally described in pebbles found in the washing concentrates of a gemstone mine in Sri Lanka (Willgallis et al., 1983). There it occurs as millimetre-size grains associated with zirconolite, baddeleyite, geikielite, spinel and perovskite. It is also reported as inclusions in pyrope crystals in mantle xenoliths from a lamprophyre pipe (Tobuk–Khatystyr field, Siberia) and from an ultramafic diatreme (Navajo volcanic field, Colorado plateau) [summary given by Wang et al. (1999)]. In the ultramafic diatreme, it occurs as small euhedral grains (5–20 µm) coexisting with rutile, ilmenite, loveringite and carmichaelite. As far as we know, srilankite has not been reported in crustal mafic granulites.
Baddeleyite and srilankite are described at the margin of ilmenite in the mafic granulites of the Lindås nappe (Fig. 2f–k). The most straightforward interpretation is that baddeleyite formed during subsolidus cooling of the magmatic intrusion, as external granules exsolved from the ilmenite. Baddeleyite was locally preserved through granulite-facies metamorphism. When associated with baddeleyite, srilankite is texturally situated between baddeleyite and ilmenite. This position suggests that it is the product of a reaction between these minerals, following the equilibrium
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The occurrence of baddeleyite and srilankite seems to be controlled by very local conditions. Except in one exceptional situation (Fig. 2d), they occur at the margin of ilmenite grains included in feldspar (or pyroxene), i.e. where no free silica was available at grain boundaries during post-magmatic cooling and metamorphism, and where ilmenite was comparatively protected from deformation during metamorphism. As srilankite occurs in various granulites and appears unrelated to the Caledonian overprint, it most probably formed during the Sveconorwegian granulite-facies event (Fig. 10).
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Zircon growth during granulite-facies metamorphism
In the granulites of the Lindås nappe, a rounded to anhedral luminescent zircon overgrowth (Th/U = 0·53; 924 ± 58 Ma) generally occurs on large zircon cores randomly distributed in the rock (Th/U = 1·27; 952 ± 32 Ma). An inner core is sometimes observed (Th/U = 0·54; 950 ± 26 Ma). ID-TIMS data unambiguously show that the overgrowth (929–933 Ma) is younger than the core (951 Ma). The oscillatory to sector zoning pattern and the high Th/U ratio of the core indicate that it is probably of magmatic origin. The equivalence of age between the inner core and the core indicate that the inner core probably formed during a first episode of saturation in the magma chamber. The higher U content (551 ppm) and lower Th/U ratio (0·54) of the inner core suggest that it may represent a pseudomorph of early crystallized baddeleyite transformed to zircon by a peritectic reaction during magmatic differentiation.
The first post-magmatic transformation recorded in these rocks is the Sveconorwegian granulite-facies penetrative metamorphism. As large anhedral zircon grains are in apparent equilibrium in the granulite-facies granoblastic texture, the luminescent anhedral overgrowth, including the ‘hat’-shaped overgrowth, is attributed to granulite-facies metamorphism. Common embayment textures at the core–overgrowth interface also attest to secondary replacement of the magmatic core by luminescent metamorphic zircon. Metamorphic zircon thus formed both by crystallization of a zircon overgrowth and by secondary replacement of magmatic zircon.
In addition, a corona of micro-zircon grains commonly occurs at the margin of ilmenite. The mere fact that the corona is located at the surface of granulite-facies ilmenite (Fig. 2a and b) suggests that the zircon corona is linked to Sveconorwegian granulite-facies metamorphism. Two supplementary arguments give support to this interpretation: (1) the grading of the zircon corona into the ‘hat’-shaped zircon overgrowth (Figs 2c and 3) with a Th/U ratio of 0·44, typical of granulite-facies zircon, makes a link between these two occurrences of metamorphic zircon; (2) the lack of Caledonian age signature at the surface of coarse-grained zircon (SIMS surface analyses in BH5 and BH10, Fig. 5) indicates that there is no significant crystallization of zircon during the Caledonian event in the granulites.
The widespread and systematic location of a zircon corona at the surface of ilmenite grains suggests (but does not demonstrate) that ilmenite was the source of its constitutive Zr. The widespread occurrence of zircon rather than baddeleyite or srilankite in the corona indicates that silica activity was high enough along grain boundaries forming an open network to stabilize zircon rather than baddeleyite or srilankite during metamorphism. Baddeleyite that was present before granulite-facies metamorphism, or that exsolved from the ilmenite during granulite-facies metamorphism, must have reacted in most places with silica available at grain boundaries to form metamorphic zircon. The texture pictured in Fig. 2d, showing a baddeleyite inclusion in ilmenite associated with a continuous zircon cover at the surface of ilmenite, may represent an exceptional situation where this reaction was frozen before completion. All in all, the textural record suggests that Zr was extracted from the ilmenite lattice as baddeleyite exsolution lamellae and that these exsolutions reacted to form zircon at the surface of ilmenite during granulite-facies metamorphism (Fig. 10).
During ductile deformation of jotunite–mangerite rocks, ilmenite can accommodate significant deformation and ilmenite grain boundaries migrate. Ilmenite can flow to pressure shadow zones, implying that ilmenite is a comparatively ductile phase at high temperature (Paludan et al., 1994; Duchesne, 1999). A corona of micro-zircons situated at the surface of an ilmenite grain that underwent major deformation during granulite-facies metamorphism is likely to have been destroyed during deformation. The micro-zircons presumably migrated along their original grain boundary, or moved away from it. They probably coalesced to form an overgrowth on large magmatic zircon grains. Formation of the ‘hat’-shaped zircon overgrowth (Fig. 3) also probably involved significant migration of small zircon grains at the surface of ilmenite and their coalescence where a magmatic zircon core was in contact with ilmenite or came into contact with ilmenite during deformation. On the contrary, the coronae of micro-zircons that are observed and preserved today at the surface of ilmenite are probably a characteristic of comparatively poorly deformed ilmenite grains.
In a wide range of granulites including the granulites of the Lindås nappe, ilmenite consists of exsolution lamellae of haemo-ilmenite and ilmeno-haematite (Fig. 2). The exsolution process typically takes place during regional cooling following granulite-facies metamorphism. The random position of the zircon grains in the corona textures relative to these exsolution lamellae suggests that the formation of the coronae and the formation of the exsolution lamellae are not related.
Zircon growth during eclogite-facies metamorphism
In the samples affected by static Caledonian overprint, the zircon corona textures are preserved. In the amphibolite BH8, they are located at the external surface of the titanite corona replacing ilmenite (Fig. 2e). In the eclogite BH2, they occur around rutile, which is the main eclogite-facies breakdown product of ilmenite (Figs 7 and 8). In these samples the zircon coronae probably outline the surface of the former granulite-facies ilmenite grain. Their preservation indicates that fluid circulation associated with the Caledonian overprint and conversion of granulite to amphibolite- or eclogite-facies assemblages did not result in any significant dissolution of zircon (Fig. 10).
Several lines of evidence support the occurrence of a specific episode of zircon precipitation during eclogitization of the mafic granulites in Holsnøy:
- in the eclogite sample BH2, the habit of zircon is specific (Figs 4h–m and 9): a characteristic prismatic oscillatory zoned overgrowth is observed and zircon commonly occurs in polycrystalline aggregates. Oscillatory zoning of zircon is typically considered as evidence for crystallization in a magmatic environment. Oscillatory zoning is nevertheless also a common feature in crystals grown in hydrothermal conditions. There is evidence for an aqueous fluid medium during static eclogite-facies metamorphism in the Lindås nappe (Austrheim, 1987) allowing for growth of oscillatory zoned metamorphic minerals, including zircon, in these eclogites. The occurrence of polycrystalline aggregates (Fig. 4h and l) indicates that nucleation of zircon occurred at several sites located close to each other and that the crystals came into contact during simultaneous growth. This feature has been described for gabbro affected by static granulite-facies metamorphism (Davidson & van Breemen, 1988). In this metagabbro, polycrystalline aggregates of columnar zircon formed radially around baddeleyite, replacing it during metamorphism. Polycrystalline zircon aggregates are thus fully compatible with a metamorphic environment.
- A Caledonian zircon overgrowth is observed only in the eclogite-facies sample. It yields an average SIMS age of 455 ± 29 Ma (Figs 5 and 6) and an ID-TIMS age of 456 ± 7 Ma. In the other samples, including the amphibolite-facies overprinted samples, SIMS analyses of sections and external surfaces of zircon failed to provide any Caledonian age signature and thus any evidence for a significant Caledonian overgrowth.
- In the eclogite, a zircon corona occurs around rutile (Fig. 8). In one example, in the thin section of BH2 (Fig. 8b), a prismatic zircon crystal (25 µm wide) is observed aligned with a corona of small rounded to flat zircon grains (<10 µm in length). CL imaging of this crystal and its surrounding matrix shows that it consists of a small core (6 µm in diameter) in perfect continuation with the corona, surrounded by a prismatic zoned overgrowth (Fig. 11). This texture suggests that a small zircon in the corona served as a nucleus for the crystallization of an euhedral to prismatic overgrowth. This is taken as evidence for zircon precipitation after the formation of the corona, namely during Caledonian eclogite-facies metamorphism. By extrapolation, this occurrence suggests that a number of prismatic crystals and polycrystalline aggregates extracted from the mineral separates, showing a small rounded to flat core surrounded by an euhedral overgrowth (Fig. 4h–m), could be derived from similar textures. The prismatic overgrowth analysed from such zircons in the mineral separates yields a Caledonian average SIMS age of 455 ± 29 Ma (Figs 5 and 6).
- The euhedral Caledonian overgrowth in the eclogite BH2 is distinctly poor in Th ( 7 ppm), resulting in a Th/U ratio 0·13. The Th/U ratio of zircon depends on several factors including its crystallochemical properties, varying with temperature, the equilibrium between the different Th- and U-bearing phases, and the whole-rock Th and U contents. In the eclogites, apatite is free of needle-shaped inclusions of monazite, but commonly associated with, or surrounded by allanitic epidote (Austrheim et al., in preparation). Clinozoisite is part of the mineral assemblage. The propensity of epidote-group minerals to act as a sink for rare earth elements and Th is well established (Deer et al., 1986). The co-precipitation of metamorphic zircon with epidote-group minerals during eclogite-facies metamorphism can account for the very low Th content and Th/U ratio of metamorphic zircon (Fig. 5).
Transformation of the granulite-facies assemblage of two pyroxenes + garnet + plagioclase + ilmenite to form the eclogite-facies assemblage of garnet + omphacite + rutile + epidote + phengite + amphibole thus resulted in release of Zr (and SiO2), and precipitation of zircon (Fig. 10). As there is evidence for fluid-present conditions during eclogite-facies metamorphism, and as the zircon overgrowth is oscillatory zoned, the Zr hosted in the granulite-facies assemblage may have been into solution during the metamorphic transition, and precipitated as zircon around existing zircon nuclei. In this situation, the spatial relationship between the zircon overgrowth and the corona of micro-zircons around rutile is not direct evidence for derivation of the Zr mainly from ilmenite. In granulite sample BH5, in which various thick eclogite-facies symplectites of garnet, omphacite and amphibole are widespread at grain boundaries, and in which large ilmenite grains are not transformed to rutile, no Caledonian zircon overgrowth can be detected with the SIMS. This suggests that only total transformation of the granulite- to the eclogite-facies assemblage, including the breakdown of ilmenite to rutile, results in precipitation of zircon. The specific role of the ilmenite to rutile breakdown reaction in releasing Zr is outlined by this argument, but cannot be demonstrated.
Formation of the amphibolite-facies assemblage, amphibole + biotite + titanite + epidote, in contrast, is not associated with any new zircon growth (samples BH11 and BH8). It has been shown that amphibole in some metabasic rocks contains a relatively high concentration of Zr (>100 ppm; Fraser et al., 1997); this indicates that Zr released during breakdown of the granulite-facies assemblage may be accommodated in amphibole, titanite or biotite.
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CONCLUSIONS |
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TOPABSTRACTINTRODUCTIONMETAMORPHISM IN THE LINDAS...METHODSGRANULITES AND AMPHIBOLITESSIMS Th-U-Pb data on...ECLOGITESIMS Th-U-Pb data on...DISCUSSIONCONCLUSIONSREFERENCES |
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In mafic granulites of the Lindås nappe, baddeleyite, srilankite and zircon micro-grains are observed at the margins of ilmenite. A discontinuous corona of zircon grains around ilmenite is a widespread feature. Baddeleyite is interpreted as an exsolution product from magmatic ilmenite, and srilankite and zircon as products of Sveconorwegian granulite-facies metamorphism. Zircon formed by reaction between baddeleyite and silica available at grain boundaries. Srilankite formed only in quartz-undersaturated subsystems, i.e. where ilmenite is surrounded by feldspar, by reaction between baddeleyite and ilmenite. Textural relationships suggest that magmatic ilmenite acted as a source (not necessarily the only source) of Zr for new growth of metamorphic zircon during granulite-facies metamorphism (Fig. 10).
In rocks affected by static amphibolite- or eclogite-facies Caledonian overprint, the corona of micro-zircons is preserved after ilmenite breakdown. The eclogite sample contains a specific zircon population with a Caledonian (455 ± 29 Ma), euhedral, oscillatory zoned, low-Th overgrowth (Th/U 0·13). A prismatic zircon overgrowth is reported around micro-zircon in a corona texture. It is concluded that the precipitation of the euhedral overgrowth is linked to the breakdown of the granulite-facies assemblage of two pyroxenes + garnet + plagioclase + ilmenite to form the eclogite-facies assemblage of garnet + omphacite + rutile and minor epidote + phengite + amphibole, in the presence of a fluid. The textural relationships do not demonstrate that ilmenite is the main source of Zr. The coeval crystallization of rutile, zircon and epidote group minerals is probably associated with a partitioning of Th into epidote. This partitioning can account for the low Th content of eclogite-facies zircon. In the amphibolite samples, SIMS analyses failed to reveal any Caledonian zircon crystallization. It is concluded that the Caledonian amphibolite-facies overprint, characterized by the formation of an amphibole + biotite + titanite + epidote assemblage, did not result in any significant dissolution or precipitation of zircon.
An episode of oscillatory zoned zircon growth coeval with fluid-driven eclogite-facies forming reactions is documented. U–Pb geochronology applied to this type of zircon thus records fluid infiltration leading to eclogitization. Variation of zircon ages across a high-P terrain may thus reflect the propagation of fluids and fluid-driven eclogitization in the crust, rather than variations of pressure–temperature conditions.
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ACKNOWLEDGEMENTS |
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B.B. benefited from an EU TMR research fellowship at the University of Oslo. SIMS data were acquired at the Nordsim laboratory, co-funded by the Norwegian Research Council. T. Sunde and J. Vestin are thanked for supervision during data acquisition. Thanks are due to G. Bye-Fjeld, E. Lieng and T. Winje for technical assistance, to E. Essene for advice in analysing srilankite, to P. Padget for reading the manuscript, and to D. Ellis, G. Fraser and F. Oberli for constructive reviews. The project was partially funded by Norwegian Research Council project 107603/410. This is Nordsim Contribution 25.
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FOOTNOTES |
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*Corresponding author. Telephone: +47-73-90-4000. Fax: +47-73-92-1620. E-mail: bernard.bingen{at}ngu.no
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REFERENCES |
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