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Journal of Petrology | Volume 45 | Number 2 | Pages 415-437 | 2004
© Oxford University Press 2004; all rights reserved
An Archean(?) to Paleozoic Evolution for a Garnet Peridotite Lens with Sub-Baltic Shield Affinity within the Seve Nappe Complex of Jämtland, Sweden, Central Scandinavian Caledonides
1 SCHOOL OF EARTH AND ENVIRONMENTAL SCIENCES, QUEENS COLLEGE AND THE GRADUATE CENTER, CUNY, FLUSHING, NY 11367, USA
2 LAMONT–DOHERTY EARTH OBSERVATORY OF COLUMBIA UNIVERSITY, PALISADES, NY 10964, USA
3 VENING MEINESZ RESEARCH SCHOOL OF GEODYNAMICS, FACULTY OF EARTH SCIENCES, UTRECHT UNIVERSITY, BUDAPESTLAAN 4, PO BOX 80.021, UTRECHT, 3508 TA, NETHERLANDS
4 ARC NATIONAL KEY CENTRE GEMOC, DEPARTMENT OF EARTH AND PLANETARY SCIENCES, MACQUARIE UNIVERSITY, SYDNEY, N.S.W. 2109, AUSTRALIA
RECEIVED NOVEMBER 25, 2002; ACCEPTED JULY 21, 2003
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ABSTRACT |
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TOP
ABSTRACT
INTRODUCTIONMineralogical, isotopic, geochemical and geochronological evidence demonstrates that the Friningen body, a garnet peridotite body containing garnet pyroxenite layers in the Seve Nappe Complex (SNC) of Northern Jämtland, Sweden, represents old, certainly Proterozoic and possibly Archean, lithosphere that became incorporated into the Caledonian tectonic edifice during crustal subduction into the mantle at c. 450 Ma. Both garnet peridotite and pyroxenite contain two (M1 and M2) generations of garnet-bearing assemblages separated by the formation of two-pyroxene, spinel symplectite around the M1 garnet and the crystallization of low-Cr spinel1C in the matrix. These textures suggest initial high-pressure (HP) crystallization of garnet peridotite and pyroxenite succeeded by decompression into the spinel stability field, followed by recompression into the garnet peridotite facies. Some pyroxenite layers appear to be characterized solely by M2 assemblages with stretched garnet as large as several centimeters. Laser ablation microprobe–inductively coupled plasma mass spectrometry Re–Os analyses of single sulfide grains generally define meaningless model ages suggesting more than one episode of Re and/or Os addition and/or loss to the body. Pentlandite grains from a single polished slab of one garnet peridotite, however, define a linear array on an Re–Os isochron diagram that, if interpreted as an errorchron, suggests an Archean melt extraction event that left behind the depleted dunite and harzburgite bodies that characterize the SNC. Refertilization of this mantle by melts associated with the development of the pyroxenite layers is indicated by enriched clinopyroxene Sr–Nd isotope ratios, and by parallel large ion lithophile-enriched trace element patterns in clinopyroxene from pyroxenite and the immediately adjacent peridotite. Clinopyroxene and whole-rock model Sm–Nd ages (TDM = 1·1–2·2 Ga) indicate that fertilization took place in Proterozoic times. Sm–Nd garnet2–clinopyroxene2–whole rock ± orthopyroxene2 mineral isochrons from three pyroxenite layers define overlapping ages of 452·1 ± 7·5 and 448 ± 13 Ma and 451 ± 43 Ma (2
). All ages are within error of Sm–Nd mineral ages from eclogite in the enclosing host gneiss, demonstrating that, whereas silicate intrusion and fertilization occurred in the mantle during the Proterozoic, the formation of at least the M2 garnet-bearing assemblages occurred in the crust, as that crust was subducted into the mantle during the Caledonian orogenic cycle. By analogy we infer a similar origin for other mantle-derived lenses in HP nappes of the SNC in Northern Jämtland. KEY WORDS: garnet peridotite; HP/UHP terranes; Scandinavian Caledonides; Seve–Koli Nappe Complex; Re–Os, Sm–Nd and Rb–Sr systems; geothermobarometry; subcontinental mantle
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INTRODUCTION |
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The Scandinavian Caledonides were formed during the closure of the Aegir and Iapetus oceans separating Baltica in the east from, respectively, Russia to the north and Laurentia in the west (Torsvik et al., 1992
, 1996). The culminating collision that closed these oceans occurred during the 420–390 Ma Scandian Orogeny when Laurentia collided with Baltica. This collision resulted in a nappe complex that was thrust to the east over the Baltoscandian basement and its autochthonous sedimentary cover. The nappes are subdivided into the Lower, Middle, Upper and Uppermost Allochthons (Fig. 1a; from Roberts & Gee, 1985). The Lower and Middle Allochthons were derived from the overridden continent, Baltica, whereas the Uppermost Allochthon contains exotic continental segments and arc complexes interpreted to represent either fragments of Laurentia or nearby arc terranes. The allochthon between the Uppermost and Middle Allochthons, somewhat confusingly called the Upper Allochthon, is actually a variety of thrust-bounded terranes derived either from the thinned outermost edge of Baltica or from more outboard, oceanic terranes (Stephens, 1988; Roberts & Stephens, 2000).
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The Upper Allochthon, containing the peridotite bodies that are the subject of this study, is called the Seve–Køli Nappe Complex (SKNC) in the Swedish Caledonides, and extends for a distance of more than 800 km along the eastern side of the mountain belt (Fig. 1a). The lower part of the SKNC, the Seve Nappe Complex (SNC), is interpreted to represent the outermost part of Baltica, which had been stretched and rifted during the opening of the Aegir sea. The upper part, called the Køli Nappe Complex (KNC), represents more outboard oceanic terranes interpreted to be largely from Iapetus (Stephens, 1988
; Stephens & Gee, 1989; Andreasson et al., 1998; Roberts & Stephens, 2000).
Two major Caledonian orogenic events (the Finnmarkian and the Scandian) have been identified in the SKNC so far. The Finnmarkian phase is generally interpreted as a collision between an island arc (the Virisen Arc) and mainland Baltica at around 500 Ma (Sturt, 1984; Mørk et al., 1988; Essex et al., 1997; Andreassson et al., 1998). The Scandian phase, as already noted, is related to final closure of Iapetus when Baltica collided with Laurentia around 420–390 Ma (Roberts & Stephens, 2000; Carswell et al., 2003).
Mantle fragments (i.e. peridotite lenses) occur within the metamorphic rocks of the SKNC, in both the KNC and the SNC. These bodies have been interpreted as basal parts of highly dismembered ophiolite complexes (Du Rietz, 1935; Zachrisson, 1969; Zwart, 1974; Stigh, 1979; Qvale & Stigh, 1985; Bucher-Nurminen, 1988, 1991; Andreasson, 1994). Most are homogeneous, tectonized harzburgite and dunite, consistent with derivation from depleted oceanic lithosphere. However, a few bodies, all within high-pressure (HP) terranes, contain garnet and/or clinopyroxene-bearing assemblages with textures and chemistries that suggest possible derivation from the subcontinental lithosphere (Van Roermund, 1989). If so, the bodies are alpine-type peridotites (orogenic lherzolites) with completely different origins and intrusion histories from ophiolites.
We present evidence that one such garnet-bearing peridotite body, in the Swedish county of Jämtland, represents old, possibly Archean, lithosphere that was intruded and metasomatized by pyroxenite-related melts during the Proterozoic. This mantle fragment can thus not be related to the relatively younger Iapetus (or Aegir) oceanic lithosphere but must instead be a fragment of an older, presumably Baltoscandian, subcontinental lithosphere. This interpretation is consistent with changes in mineral assemblages and mineral chemistries that suggest complicated metamorphic histories for the peridotites. It is also consistent with relatively ‘enriched’ Sr–Nd isotope ratios and trace element chemistries of clinopyroxene separates, with Proterozoic Sm–Nd clinopyroxene model ages, and with possible Archean or early Proterozoic Re–Os errorchrons from sulfides. Sm–Nd mineral ages from garnet-bearing assemblages suggest that the tectonic incorporation of the peridotite fragments into the crust did not take place until the Caledonian orogenic cycle. Surprisingly, this event did not take place during the well-known Finnmarkian or Scandian Orogenies, but rather during an as yet unnamed episode of crustal subduction (the Jämtland Orogeny?) into the mantle at around 450 Ma (Brueckner & Van Roermund, 2003; Van Roermund & Brueckner, in prep.).
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REGIONAL GEOLOGY |
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The Seve Nappe Complex in North Jämtland (Tången–Inviken area; Fig. 1b) consists of medium- and high-grade schist, gneiss, quartzite and meta-arkose enclosing subordinate marble, together with extensive metabasite, and scattered ultramafic bodies (Van Roermund & Bakker, 1984
). Biostratigraphic evidence is lacking, but geochronological data suggest that at least some of the paragneiss was derived from (sedimentary or crystalline) Middle Proterozoic protoliths (Reymer et al., 1980; Claesson, 1982, 1987; Mearns & Van Roermund, 1985). The quartzite and meta-arkose are intruded by metabasic sills and dikes, an association that suggests an origin as rift-related late Proterozoic–early Phanerozoic ‘sparagmite’ intruded by diabase dikes during the opening of the Aegir Sea.
The SNC in the Tången–Inviken area is subdivided into medium-pressure and eclogite-bearing high-pressure thrust sheets [Fig. 1b and c, after Van Roermund & Bakker (1984) and Van Roermund (1985, 1989)]. Garnet peridotite occurs exclusively in the HP tectonic units, suggesting an intimate relationship between garnet peridotite and HP rocks. So far only two garnet-bearing peridotite lenses have been located in the Tången–Inviken area; one near Lake Friningen and another along the northern shore of Lake Store Jougdan (Fig. 1b, from Van Roermund, 1989). Other garnet-bearing bodies presumably exist but have not yet been located, as only a few of the hundreds of bodies that exist in the region were examined carefully for garnet peridotite or pyroxenite. Some of these bodies contain lower-grade assemblages (primarily amphibolite) that originally may have been garnet-bearing rocks. This study focuses on the Lake Friningen peridotite–pyroxenite body (GPS coordinates: 64°44'435''N, 014°33'442''E), which is exposed close to the structural base of an eclogite-bearing migmatitic kyanite–sillimanite–K-feldspar gneiss (Ertsekey lens) at the NE margin of Lake Friningen (Fig. 1b).
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FIELD OBSERVATIONS |
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The c. 200 m x 30 m peridotite body near Lake Friningen (Fig. 1b) was initially described by Van Roermund (1989)
. Much of the body is exposed along a relatively inaccessible cliff and so the body has not been mapped in detail. The peridotite is locally serpentinized, but most of it is remarkably fresh, garnet-free dunite and harzburgite. However, garnet lherzolite was found near its upper contact and at the southeastern side of the body near the lower contact. More garnet lherzolite occurs in boulders below the body. A conspicuous 0·3 m thick, >50 m long garnet pyroxenite dike cuts the lens. Thin (centimeter-scale) garnet pyroxenite layers occur throughout the massif and within some of the boulders. Within the boulders garnet lherzolite is commonly associated with these thin pyroxenites. The garnet peridotite body displays a compositional layering (S0), defined by variable amounts of the primary rock-forming minerals olivine, clinopyroxene, orthopyroxene and garnet. However, the most conspicuous feature of the garnet lherzolite is its pronounced porphyroclastic texture. Elongated, coarse-grained (up to 7 cm), heavily strained, porphyroclasts of olivine, clinopyroxene, orthopyroxene, garnet and spinel float in a mosaic to tabular recrystallized, strain-free, matrix composed of fine-grained garnet, spinel, olivine, clinopyroxene, orthopyroxene ± amphibole, defining the dominant foliation (S1). S1 is always subparallel to S0 as well as to the pyroxenite dikes. Subordinate late shear zones (S2) transect the main foliation locally. Some of the samples analyzed in this study were from the original collection of Van Roermund (1989); others were collected from the top of the body and from talus boulders beneath the body during a 2 day expedition by two of the authors (H.v.R. and H.B.) accompanied by D. A. Carswell in 2000. The samples are described below. Analytical methods and techniques used in studying the samples are presented in the Appendix. |
MICROSTRUCTURES, MINERAL CHEMISTRIES AND P–T CONDITIONS |
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The primary mineral assemblage (M1A) consists of variable amounts of the coarse-grained (e.g. 0·5–7·0 cm sized garnet) minerals garnet1A–olivine1A–orthopyroxene1A–clinopyroxene1A ± Fe–Ni-sulfides. Two types of M1A textures were recognized: coarse-grained equigranular and porphyroclastic (Nicolas & Poirier, 1976
). The equigranular type consist of coarse-grained, heavily strained and fractured M1A minerals with slightly curved grain boundaries and numerous triple junctions. Recrystallized, strain-free M2 minerals are absent or rare. The porphyroclastic M1A type shows relict M1A minerals forming heavily strained porphyroclasts within a fine-grained, strain-free, M2 matrix. The matrix minerals inhibit identification of the original M1A grain boundary geometry. Coarse-grained equigranular M1A textures are relatively scarce within the body, but do occur within some garnet pyroxenite–websterite dikes, in strong contrast to the porphyroclastic M1A textures, which are widespread. Representative electron microprobe (EMP) analyses of M1A minerals are given in Table 1. EMP line scans across M1A minerals are uniformly flat except for the outermost rims. In addition, M1A pyroxenes show pronounced exsolution textures (Fig. 2a) with spinel, garnet and pyroxene as exsolved phases. Between these exsolved phases the host minerals continue to reveal flat chemical profiles. P–T estimates were calculated from the mineral chemistries of the host minerals within these flat areas.
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M1A garnet contains Cr2O3 up to 2·7 wt % and has Mg number = 79–81. M1A olivine has Mg number = 91·5–92·5. M1A orthopyroxene contains 2·0–3·0 wt % Al2O3 and has Mg number = 91–92. M1A clinopyroxene contains 1·8–2·0 wt % Na2O and up to 5·6 wt % Al2O3, and has Mg number = 95. Based on their Ca, Fe and Mg content, both M1A pyroxenes would be classified as ‘low’-T pyroxenes (Table 1). It should be emphasized, however, that no attempt was made to recalculate primary M1A pyroxene compositions by combining their current mineral compositions (Table 1) with the relative volume percentages of the exsolved mineral phases (Fig. 2a) and their compositions. Such an attempt would result in much higher T estimates (for M1A pyroxenes) than those obtained by geothermometry using their current mineral chemistry (see below).
Application of the pyroxene–garnet geothermometers of Ellis & Green (1979), Harley (1984), Powell (1985) and Krogh (1988), and the geobarometer of Brey & Kohler (1990) to M1A mineral compositions of porphyroclastic garnet lherzolite gives T = 690–800°C and P = 11–16 kbar (Table 1) with TKrogh being around 100°C lower than TEG, TH and TP (taking all Fe as FeO). Single- or two-pyroxene thermometers indicate a similar T range. Application of the olivine–garnet geothermometer of O'Neill & Wood (1979) and O'Neill (1980) with the geobarometer of Brey & Kohler (1990) gives T = 680–700°C and P = 11–12 kbar. It should be noted that the majority of these M1A P–T calculations plot outside the stability field of garnet peridotite (at the lower-pressure side). Either the P–T estimates do not represent meaningful geological conditions, or they imply that garnet within lherzolite existed under metastable conditions.
When the geothermometers of Ellis & Green (1979) and Harley (1984), and the geobarometer of Brey & Kohler (1990) are applied to garnet pyroxenite with coarse-grained equigranular M1A textures they give T = 750–800°C at P = 14–15 kbar (Table 1). Although these calculated P–T conditions also plot below the garnet peridotite stability field, they do fall within the stability field of garnet pyroxenite.
In peridotite most M1A garnet is partially to completely overprinted by either M1B or M2 assemblages. The M1B assemblage consists of a spinel1B–orthopyroxene1B–clinopyroxene1B symplectite with simultaneous growth of coarse, reddish spinel (spinel1C; Cr number = 0·2) in the olivine matrix (Figs 2b and 3a and b). Our interpretation that the coarse, reddish spinel is part of the M1B assemblage is based on microstructural relationships in combination with the low Cr number of the spinel1C. However, an alternative interpretation for the coarse, reddish spinel1C is that it formed part of the M1A assemblage (and therefore would be termed spinel1A). The spinel-bearing M1B assemblage indicates decompression either during the late stages of M1A recrystallization or during a discrete later event.
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The M2 assemblage indicates subsequent recompression during a second episode of HP metamorphism, which converted the M1A and M1B assemblages into a fine-grained garnet2–spinel2–olivine2–orthopyroxene2 ± amphibole assemblage (Fig. 3). M2 metamorphism formed clear garnet (garnet2) rims around the coarse-grained matrix spinel (spinel1C) (Fig. 3b), caused concomitant compositional zoning in spinel1C (an increase in Cr number from 20 to 50 towards the contact with garnet2; Table 1, Fig. 3b) and formed small, discrete, new garnet2 grains within the two-pyroxene + spinel1B symplectite. In other cases M2 metamorphism fully recrystallized the M1B two-pyroxene + spinel symplectites into garnet2–spinel2 (Cr number = 0·5)–olivine2–orthopyroxene2 ± amphibole assemblages. Similar M2 coronas are also present around relicts of primary garnet (garnet1A) in the lherzolite (Fig. 3a). In these cases the relict M1A garnet cores (sample Fr-1A1, Table 1) reveal clear microfractures filled by M2 mineral assemblages including garnet2 and spinel2 (Fig. 3c). The microfractures do not continue into the M2 corona assemblage. Representative M2 mineral compositions are given in Table 1.
M2 garnets in lherzolite and some pyroxenite layers are riddled with tiny spinel grains. The garnet can be compositionally zoned with more Mg-rich, Fe-poor compositions at their margins (Van Roermund, 1989). In contrast, core to rim changes in the composition of the tiny M2 spinel grains have not been detected. M2 garnet (in lherzolite) is Cr-poor compared with M1A garnet (Table 1).
Orthopyroxene2 compositions range between 0·85 and 1·50 wt % Al2O3, with the most dominant values being close to 1·1–1·2 wt %. These values are distinctly lower than those in orthopyroxene1A, which range from 2·0 to 3·0 wt %. In addition, M1A orthopyroxene clasts in lherzolite reveal high flat Al2O3 (2·0–3·0 wt %) core compositions (see above), which significantly decrease towards flat 0·8–1·2 wt % Al2O3 values at their rims. This decrease is consistent with our interpretation that pressure increased during M2 recrystallization.
Application of the pyroxene–garnet geothermometers of Ellis & Green (1979), Harley (1984), Powell (1985) and Krogh (1988), and the geobarometer of Brey & Kohler (1990) to M2 mineral assemblages give T = 700–800°C and P = 20–30 kbar, indicating that the M2 metamorphic event occurred at significantly higher pressures than M1 (Table 1). However, similar P–T conditions can be calculated using the rim compositions of M1A porphyroclasts, which would be consistent with a single prograde event during the development of the M1A and M2 assemblages, but would not explain the development of the spinel-bearing M1B symplectites around the M1A porphyroclasts.
Application of the olivine–garnet thermometer of O'Neill & Wood (1979) and O'Neill (1980) with the geobarometer of Brey & Kohler (1990) to M2 mineral assemblages gives T = 450–600°C and P = 11–16 kbar. These P–T estimates are, however, ignored as the temperature range defined by TO'Neill plots well below the minimum temperature required to form the migmatites that characterize the surrounding gneisses (Van Roermund, 1985, 1989).
Microstructures in pyroxenite and peridotite are similar, except for the lack of olivine in the pyroxenite (compare Figs 2c and 3a). The most widespread garnet and pyroxene microstructure in the pyroxenite consists of dynamically recrystallized M2 grains [see fig. 3B of Van Roermund (1989)]. However, some of the thicker garnet pyroxenite layers preserve clear porphyroclastic garnet microstructures with garnet cores being surrounded by fine-grained garnet rims (Fig. 2c). A compositional distinction between cores and rims could not be determined and it is not certain that the cores and rims represent garnet1A and garnet2, respectively. They may represent zoning changes caused by changes in P–T during a single metamorphic progression. Like the M1A garnet microstructures in lherzolite (Fig. 3a), these porphyroclastic garnet cores in pyroxenite are heavily fractured and do not continue into the recrystallized rims (Fig. 2c). Associated matrix pyroxene is mostly recrystallized into the finer-grained M2 assemblage although some pyroxene porphyroclasts were still identified within the thickest garnet pyroxenite layer that may be M1. In contrast to M2 matrix pyroxene, these pyroxene porphyroclasts are always heavily strained and reveal clear exsolution microstructures involving garnet, spinel and pyroxene.
Application of the Ca–Cr test of Brenker & Brey (1997) to single M1A and M2 garnet compositions from lherzolite indicates that both garnet compositions plot on the same isoline but at distinctly higher temperatures and pressures than recorded by the above-quoted geothermometers. This discrepancy probably indicates that M1 garnet formed at much higher temperatures than indicated by the Fe–Mg and Ca thermometers discussed above. Such an interpretation is also consistent with the inferred higher temperatures deduced from the pyroxene exsolution microstructures. In contrast, the higher temperatures indicated for M2 garnets can be ignored because the presence of amphibole2 in the M2 assemblage makes the application of the Ca–Cr test to these rocks useless.
In summary, P–T conditions during the development of the M1A mineral assemblage could not be derived by applying standard thermobarometric techniques to the M1A mineral compositions. Indirect evidence, however, indicates that the M1A mineral assemblages formed within the garnet peridotite stability field at temperatures >800°C. These inferred M1A P–T conditions were followed by decompression towards P–T conditions that led to the development of the M1B symplectite assemblage. This decompression was also responsible for the changes in the M1A mineral compositions that led to the lower P–T estimates in the equigranular (750–800°C and 14–15 kbar) and porphyroclastic (690–800°C and 11–16 kbar) M1A assemblages, implying that the M1A garnet was metastable during this decompression. Subsequent recompression occurred during M2 (700–800°C and 20–30 kbar).
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SAMPLE DESCRIPTIONS |
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Clinopyroxene grains were separated from several garnet peridotite and pyroxenite localities for major and trace element and isotope analyses. Representative modes of the samples analyzed are presented in Table 2. Most of the clinopyroxene is believed to be M2 unless otherwise noted. 83-Fr0, Fr00-12 and Fr00-13 were selected for Sm–Nd mineral dating. Sample 83-Fr0, collected from a large dike during an earlier study by Van Roermund (1989)
, is a massive garnet websterite with very pale red, 2–4 mm garnet and pale green clinopyroxene. The garnet is riddled with tiny spinel grains (<0·1 mm); this feature suggests that the garnet and associated minerals are M2. Sample Fr00-12 is a amphibole-bearing olivine websterite from the boulder field beneath the body that contains very dark red, coarse garnet grains, some of which are 5 cm in diameter. Closer inspection shows the garnet grains to be mosaics containing veinlets and islands of millimeter-sized clinopyroxene, orthopyroxene and amphibole. The matrix between the large garnet grains is largely composed of medium- to coarse-grained clinopyroxene, orthopyroxene and amphibole. The large size of the garnet grains suggests they are M1, but they are riddled with tiny spinel grains, a feature that suggests they are M2 garnet.
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Sample Fr00-13 is one of the large garnet megacrysts, associated with minor clinopyroxene, orthopyroxene and amphibole, from the pyroxenite layer represented by Fr00-12. The garnet is clearly zoned with dark red cores and paler rims, consistent with an increase in MgO towards the rims. The garnet from this sample was separated into light and dark fractions to see if they would give significantly different (i.e. M1 and M2) ages.
Sulfides, primarily pentlandite, but also some pyrrhotite and chalcopyrite, were found in three peridotite samples from the Friningen body (Fr-8, Fr00-8 and Fr00-19). Fr-8 was collected from the boulder field and contains centimeter-sized garnet grains. The garnet is riddled with spinel grains and contains isolated multiphase inclusions of olivine, orthopyroxene, amphibole and sulfides. Fr00-8 is a peridotite immediately adjacent to pyroxenite Fr00-12 and is composed of coarse garnet in a finer-grained, equigranular, matrix of olivine–clinopyroxene–orthopyroxene–amphibole and sulfides. Fr00-19 is a porphyroclastic garnet peridotite from the top of the Friningen body with coarse olivine, garnet and clinopyroxene in a matrix of finer olivine and rare clinopyroxene. The large garnet and clinopyroxene grains are largely converted to (M1B) symplectites, suggesting that they and the associated large olivine grains are probably M1A. All sulfides observed in thin section occur within the matrix minerals, with the exception of one that occurs within an olivine porphyroclast. However, they may have originated within M1 minerals and are now relicts that survived the recrystallization of the host mineral into M2 assemblages.
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TRACE ELEMENT AND ISOTOPE GEOCHEMISTRY |
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Major and trace element data from Friningen M2 clinopyroxene, amphibole and garnet are tabulated in Table 3. The clinopyroxene grains exhibit trace element patterns on chondrite-normalized trace element and rare earth element (REE) diagrams (Fig. 4a and b) that would not be expected of clinopyroxene from the melt-depleted suboceanic lithosphere, represented, presumably, by abyssal peridotite. Friningen clinopyroxene has relatively undepleted Ti/Zr ratios (73–232), strong light REE (LREE) enrichments (Ce/Yb = 7·8–169), and relatively high Sr contents (23–224 ppm), which contrast markedly with the depleted Ti/Zr ratios (c. 250 to c. 4000), LREE-depleted patterns (Ce/Yb ratios are usually between 0·002 and 0·1) and low Sr contents (0·05–10 ppm) of most abyssal clinopyroxenes (Johnson et al., 1990
; Snow et al., 1994; Hellebrand et al., 2001, 2002; Salters & Dick, 2002). Direct comparison is complicated for some elements, such as the heavy REE (HREE), Zr and Ti, by the presence of garnet within the Friningen samples, as clinopyroxene from abyssal peridotite coexists with spinel. However, even when the elements that occur in appreciable concentrations within garnet are added to the clinopyroxene, the basic enriched pattern for Friningen clinopyroxene remains. For example, REE patterns for garnet and clinopyroxene are shown (Fig. 4b) for sample Fr00-15 (garnet pyroxenite; 45% garnet, 45% clinopyroxene; 10% orthopyroxene). When the REE in the garnet are added to the clinopyroxene (dashed line in Fig. 4b), it continues to show an enriched LREE pattern (Ce/Yb = 5) that contrasts sharply with the steep drop-off in Nd and the lighter REE shown by abyssal clinopyroxene.
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If trace element data for Friningen clinopyroxenes are plotted on other discrimination diagrams, such as those used by Rivalenti et al. (1996)
, they generally plot well away from the fields defined by abyssal pyroxene, but at the same time they tend to plot in areas defined by metasomatized mantle-derived xenoliths. Normalized trace element and REE patterns very similar to those shown in Fig. 4a and b are shown by clinopyroxene from xenoliths and orogenic lherzolites that have been metasomatized by a variety of fluids (O'Reilly & Griffin, 1988; Hauri et al., 1993; Rudnick et al., 1993; Norman, 1998; Rampone & Morten, 2001). The effect of metasomatism on the geochemistry of clinopyroxene from peridotite was addressed by Rivalenti et al. (1996), who concluded that (1) it is difficult to use clinopyroxene trace element data to distinguish subcontinental lithosphere from oceanic lithosphere, because (2) their trace element chemistry probably reflects mantle metasomatic processes rather than primary compositional features. We conclude that the enriched character of the Friningen clinopyroxene is not a function of initial composition, but rather of contamination by fluids enriched in large ion lithophile elements (LILE).
Rivalenti et al. (1996) also suggested that ‘clinopyroxene geochemistry does not allow a clear distinction between metasomatic agents’; that is, hydrous fluids, carbonatite melts, alkali silicate melts or combinations thereof can all yield similar trace element patterns. We might add that it is difficult to distinguish between fluids that flux within the mantle (some of which may have crustal signatures) from those that emanate directly from the crust. For example, Rampone & Morten (2001) presented trace element signatures in clinopyroxene from the garnet peridotite of the Ulten Zone of the Eastern Alps that are very similar to those shown in Fig. 4a and b, and proposed that they resulted from metasomatism by fluids derived from the surrounding country rock after the peridotites had been introduced into the crust.
Field evidence suggests two possible sources for the fluids that metasomatized the Friningen body. The peridotite within boulders at the base of the body is intimately associated with pyroxenite layers and it is reasonable to propose that the melts that led to the formation of these layers metasomatized the adjacent peridotite. Support for this model comes from the parallel patterns shown by Fr00-8, a garnet peridotite, and Fr00-11 and Fr00-12, which are from an immediately adjacent pyroxenite layer. Similar relationships have been observed and interpreted in detail elsewhere [see review by Rivalenti et al. (1996) and references therein]. However, the garnet peridotite at the top of the body (samples Fr00-19 and Fr00-20) was not observed to be associated with pyroxenite. Access to this part of the body is difficult because of the steep terrane and we cannot exclude the possibility that pyroxenite exists nearby. However, the proximity of the garnet peridotite to the upper contact of the body raises the possibility that the fluids that fertilized the peridotite could have come from the immediately adjacent, highly migmatized, gneiss, an association very similar to that described by Rampone & Morten (2001) in the Ulten terrane. The trace element patterns shown by the two samples from the top of the body are the most enriched of all samples analyzed (Fig. 4a), but in most other respects they mimic the troughs and peaks of the clinopyroxene from associated peridotite and pyroxenite. It would be discouraging if such similar patterns could result from two very different agents of metasomatism.
We provisionally interpret all patterns as largely the result of intrusion related to the formation of the pyroxenite. The field relationships are very clear within the boulders at the base of the body, and the absence of pyroxenite near the top is not unequivocally demonstrated. Model age data, presented below, indicate that the enrichment event occurred during the Proterozoic, which would be inconsistent with crustal contamination after emplacement during the Paleozoic Caledonian Orogeny. Finally, the patterns from the Friningen clinopyroxene are similar to those (shown as the shaded area in Fig. 4a) of clinopyroxene in peridotite and pyroxenite of the Western Gneiss Region (WGR) of Norway (Brueckner & Medaris, 1998). The WGR peridotite is widely regarded as representative of the Proterozoic sub-continental mantle (Jamtveit et al., 1991; Brueckner et al., 2002). The trace element patterns of clinopyroxene in WGR peridotite bodies are attributed to metasomatism related to the intrusion of melts that led to the formation of pyroxenite layers (Brueckner & Medaris, 1998). A possible link between peridotite and crustal contamination can be excluded for most WGR peridotite bodies because metasomatism occurred there before the formation of mid-Proterozoic garnet-bearing assemblages, which occurred while the peridotite bodies were still part of the sub-Baltic Shield mantle [however, see Brueckner et al. (2002) for local exceptions].
Although the evidence suggests that most metasomatism was related to the intrusion of melts that led to the formation of pyroxenite layers, there are Sr and Nd isotope data, also discussed further below, that suggest there was some metasomatism caused by hydrous fluids that emanated from the surrounding gneiss. This metasomatism was associated with the development of secondary (M2?) amphibole in some assemblages, particularly in the boulders at the base of the body. Amphibole patterns are largely similar to those shown by the clinopyroxene (Fig. 4a), indicating that the metasomatism introduced mostly LILE (Sr, Ba and the LREE), which caused apparent deepening of the high field strength element (HFSE) troughs on the normalized trace element patterns by raising the concentrations of the elements on either side of the troughs.
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GEOCHRONOLOGY AND ISOTOPE GEOCHEMISTRY |
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Sm–Nd mineral ages
Sr and Sm–Nd isotope data from Friningen clinopyroxene and other phases are presented in Table 4 and plotted in Figs 5 and 6. Two pyroxenite dikes (83-Fr0 and Fr00-12) give identical (i.e. within error) Sm–Nd garnet–clinopyroxene–whole-rock ± orthopyroxene dates (Fig. 5a and b) of 452·1 ± 7·5 Ma and 448 ± 13 Ma (all errors are 2
). As noted above, the garnet in both rock assemblages contains numerous spinel inclusions, a feature that is characteristic of M2 garnets (Van Roermund, 1989). The ages are believed, therefore, to date M2 metamorphism. M1 metamorphism is not dated. The third sample, the garnet megacryst, Fr00-13, defines an identical age but with a larger error (451 ± 43 Ma, MSWD = 3·4), which is due to the different positions of the garnet core and garnet rim on the isochron diagram (Fig. 5c). The core appears to define an older age than the rim. However, a simple clinopyroxene–garnet core regression yields an age of 448 ± 18 Ma that is statistically indistinguishable from the clinopyroxene–garnet rim age of 432 ± 18 Ma. If the core represents M1 garnet and the rim M2 garnet, there is no significant age difference between the formation of the two assemblages.
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The c. 450 Ma age defined by the pyroxenite dikes is identical, within error, to Sm–Nd mineral ages from two eclogite boudins in the nearby crustal rocks (Brueckner & Van Roermund, 2001
; Van Roermund & Brueckner, in prep.) indicating that the formation of garnet-bearing assemblages occurred at the same time in both the peridotite and the host country rock. We propose, therefore, that HP M2 metamorphism occurred as the result of crustal subduction into the mantle and that the peridotites had been introduced into the crust during or before subduction rather than after subduction had ceased. This interpretation is consistent with the zoning patterns and textures described above. The Friningen peridotite would therefore be of the ‘prograde type’ using the classification system proposed by Brueckner & Medaris (1999). The c. 450 Ma age is significantly younger than the well-established Finnmarkian Orogeny at c. 500 Ma, and significantly older than the classic Scandian Orogeny at 420–390 Ma. It implies that another major episode of crustal subduction and HP metamorphism must be invoked to explain the evolution of the Scandinavian Caledonides (Brueckner & Van Roermund, 2003). Sm–Nd model ages calculated relative to depleted mantle (TDM) from Friningen clinopyroxene grains and, where available, whole-rock analyses, range between 1·1 and 2·2 Ga for both pyroxenite and peridotite, with most ages clustered between 1·4 and 1·8 Ga (Table 4). It might be argued that determining model ages from clinopyroxene is unjustified because garnet also contains REE characterized by high Sm/Nd ratios that therefore fractionate the clinopyroxene Sm/Nd ratio away from that of the bulk sample. Although strictly true, clinopyroxene usually contains an order of magnitude more Sm and Nd than garnet and therefore the Sm/Nd ratio of the clinopyroxene is usually very close to that of the whole rock except in unusual circumstances. This similarity can be seen in Fig. 5a and b, where the whole rock and clinopyroxene plot next to each other and give nearly identical model ages (Table 4), whereas sample Fr00-13, shown in Fig. 5c, reflects an unusual situation where the garnet megacryst makes up more than 80% of the sample (Table 2) and the Sm/Nd ratio of the whole rock is very different from that of the clinopyroxene, with the result that the model ages of the two are significantly different. The advantage of using clinopyroxene instead of the whole rock occurs when the rock contains secondary phases such as amphibole, which could indicate open-system behavior and the possible late introduction of REE. In these circumstances the clinopyroxene would be more likely to give a reliable ratio.
The model ages from the Friningen body overlap equally scattered, but also mostly Proterozoic model ages (TDM = 1·5–2·1 Ga) from clinopyroxene from WGR peridotite bodies (Jamtveit et al., 1991; H. K. Brueckner, unpublished data, 2003). The WGR model ages are interpreted to reflect the time of pyroxenite-related fertilization of the sub-Baltic lithosphere where garnet-bearing assemblages either formed directly by crystallization from the melts or formed shortly thereafter through a mantle recrystallization event (Jamtveit et al., 1991; Brueckner & Medaris, 1998). The model ages from the Friningen body are also interpreted as dating the time of mantle fertilization of what previously had been depleted dunite and harzburgite. Unlike the peridotite bodies in the WGR, the fertilized Friningen body may not have been converted to garnet-bearing assemblages until about 450 Myr ago, nearly one billion years after the fertilization event. However, it should be noted that we have not obtained a reliable age from M1 garnet, which one of us (H.v.R.) believes could have formed during the Proterozoic, in which case the history of the Friningen peridotite would be identical to those of the WGR.
TCHUR model ages from clinopyroxene are significantly younger than the TDM model ages (Table 4). If the pyroxenite-related melts that caused the metasomatism were derived from a more enriched source, the TCHUR ages would more accurately date this event at between 0·8 and 1·2 Ga. Metasomatism of SNC peridotites at c. 1·0 Ga could be related to early continental break-up with associated dike intrusions along the Baltoscandian margin. This event is dated by an Sm–Nd isochron from a dolerite dike at c. 930 Ma, and was presumably characterized by high heat flow in the underlying mantle related to the throughput of magmas from the asthenosphere as the overlying lithosphere was pulled apart (Solyom et al., 1984, 1992; Johansson & Johansson, 1990; Andreasson, 1994).
Sr–Nd patterns
The present-day Sr–Nd isotope ratios of Friningen clinopyroxenes (Table 4) tend to cluster around somewhat enriched 87Sr/86Sr ratios (0·70341–0·70556) and strongly enriched (i.e. low) 143Nd/144Nd ratios (0·51165–0·51252, or 
Nd of -19 to -2·3) which takes them below the mantle trend and towards the EM I component of Zindler & Hart (1986). The clinopyroxene analyses do not show an obvious subduction zone signature characterized by large increases in 87Sr/86Sr, such as shown by clinopyroxene from peridotite bodies in the Bohemian Massif of the Variscan belt (Brueckner & Medaris, 1998). They also do not show the sub-continental mantle trend displayed by present-day ratios of clinopyroxene from most peridotite bodies in the WGR [see fig. 7A of Brueckner & Medaris (1998)], which show large variations in 143Nd/144Nd ratios but little change in very low values of 87Sr/86Sr. However, this lack of a marked trend for clinopyroxene from the Friningen body may reflect the fact that all analyses are restricted to a single body whereas those from the Bohemian Massif and the WGR are from several bodies that collectively define their distinct trends (Brueckner & Medaris, 1998).
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), Fr00-8 (
) and Fr00-19 (•). The best-fit line is for sample Fr00-19. (b). Sulfides from Fr00-19 that contain >3 ppm Os and have 187Re/188Os ratios <3·2. Samples shown by error bars are from a single thick section of sample Fr00-19 (six points) and define the ‘age’ and ‘error’ shown on the diagram.Most clinopyroxene analyses define Proterozoic Sm–Nd model ages, suggesting that their 147Sm/ 144Nd and 143Nd/144Nd ratios have not been significantly perturbed since the Proterozoic. Figure 6 is a plot of the initial 143Nd/144Nd ratios of the clinopyroxene samples based on their measured 147Sm/144Nd ratios and calculated back to 1·53 Ga (the average TDM model age) versus 87Sr/86Sr ratios also calculated back to 1·53 Ga based on a common 87Rb/86Sr ratio in clinopyroxenes of 0·01 (Rb concentrations are usually below detection limits in Table 2 and are too low to change the Sr ratios significantly). The Friningen clinopyroxene analyses have higher age-corrected 87Sr/86Sr ratios than the Proterozoic clinopyroxene from peridotite and pyroxenite bodies from the WGR, shown by the shaded fields in Fig. 6 (Brueckner & Medaris, 1998
; H. K. Brueckner, unpublished data, 2003). Their enriched 87Sr/86Sr signatures, relative to most WGR peridotites, indicate some radiogenic Sr contamination related either to subduction-related fluids while the peridotite was still in the mantle or to fluids from the host gneiss after the peridotite had been introduced into the crust. Amphibole separated from another peridotite body in the area has very high age-corrected 87Sr/86Sr ratios (Fig. 6). The amphibole formed by hydration of what probably were garnet pyroxenite veins during a relatively late stage of retrograde metamorphism when the peridotite bodies were almost certainly in the crust. The very high 87Sr/86Sr ratios of the amphiboles (Fig. 6) suggest that the fluids that caused hydration were derived from the surrounding, relatively radiogenic, crustal rocks. It is possible that the clinopyroxenes within the Friningen peridotite were contaminated, although to a much lesser extent, by similar crustal fluids. Recent work in the WGR (Brueckner et al., 2002) indicates that some peridotite bodies in the extreme NW corner of the WGR underwent similar crustal contamination during Caledonian subduction. If correct, the isotope and trace element patterns of Friningen garnet pyroxenite and peridotite are the result of two metasomatic episodes; one related to the intrusion of melts in the mantle during the mid-Proterozoic and the other related to penetration by hydrous fluids from the surrounding crust during Caledonian metamorphism.
Re–Os
Re–Os data from sulfide grains within three peridotite samples are presented in Table 5 and plotted in Fig. 7. Most sulfide grains contain measurable Re and Os concentrations. The data are most reliable when the 187Re/188Os ratio is <3·0 because the correction for interference by 187Re becomes significant at higher values. The majority of sulfide analyses were obtained from sulfides found in the non-magnetic fraction of crushed samples, which were mounted in epoxy for analyses. However, Os-bearing sulfides, specifically pentlandite grains, were also measured in situ on the surface of a polished 100 µm ‘thick’ section of sample Fr00-19. Most in situ sulfides occupy interstitial positions in the recrystallized M2 matrix so that their timing relationships relative to major phases could not be determined. One sulfide occurred within an M1 olivine porphyroclast, suggesting that its formation preceded or accompanied the formation of the M1 assemblage. The interstitial sulfides could also have been inclusions within M1 minerals, but ended up in the groundmass when the M1 minerals recrystallized into M1B or M2 assemblages. Minimum model ages (TRD) assuming that Re addition occurred either at 450 Ma or more recently (i.e. today) yield meaningless future ages in almost all cases. Model ages (TMA in Table 5) that assume Re was present at the time of sulfide formation are also scattered and meaningless for most samples. Some sulfides from Fr00-8 are composites of either pentlandite and pyrrhotite or pentlandite and chalcopyrite. It is possible for both chalcopyrite and pentlandite to form from a single sulfide melt (Alard et al., 2000). However, the composite textures may also indicate that sample Fr00-8 underwent two or more sulfide-forming events, a relatively common phenomenon in mantle peridotite (Alard et al., 2000; Pearson et al., 2002). Multiple sulfide-forming events could be responsible for the widely varying 187Os/188Os and Re/Os ratios.
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In contrast, sulfides from sample Fr00-19 appear to have formed through a single event. Neither chalcopyrite nor pyrrhotite was found associated with the pentlandite grains in a thin section of sample Fr00-19. When the pentlandite analyses from the sample are plotted on an Re–Os diagram, they generate an array that scatters around a best-fit line (Fig. 7a, samples with 187Re/188Os > 3·0 are excluded) with an ‘age’ of 2665 ± 550 Ma (MSWD = 52). If only samples with >3 ppm Os are considered, the array becomes more linear (Fig. 7b) and yields an ‘age’ of 2880 ± 350 Ma (MSWD = 33). If only the sulfides that were measured in situ on a single thick section from sample Fr00-19 are considered (the six points with error bars in Fig. 7b), the array nearly becomes an isochron with an age of 2970 ± 470 Ma (MSWD = 2·4) and an initial 187Os/188Os ratio of 0·1182. A nearly identical linear array with a similar c. 3·0 Ga apparent age was obtained for a microdiamond-bearing megacrystic orthopyroxenite in the WGR of Norway (Brueckner et al., 2002
). Arguably, sample Fr00-19, or the micro-domain represented by the analyzed thick section of this sample, did not suffer a later generation of sulfide formation, or perhaps did so to a lesser extent than sample Fr00-8, and so preserved the pre-middle Proterozoic history of this peridotite body. If so, the age suggests an Archean origin and, if the sulfides are interpreted as residues of melting, the Archean age may date the melt extraction event that depleted the Archean mantle to produce the depleted dunite and harzburgite bodies that characterize most parts of the Friningen body, and most of the other peridotite bodies in the HP nappes of the Jämtland region. Alternatively, the array could be a meaningless mixing line. The array defines a rather high initial 187Os/188Os ratio of 0·118, which is significantly higher than the commonly accepted value of primitive mantle at 3·0 Ga of 0·106. If the array has age significance it would imply that the mantle already had a ‘crustal’ component during Archean melting or that the sulfides were derived from a more evolved source and were introduced into the peridotite rather than being refractory residues.
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CONCLUSIONS |
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The Friningen body potentially records an evolution spanning from 0·45 Ga to possibly c. 3 Ga. Textural evidence and mineral chemistry indicate two episodes of HP metamorphism, within the garnet stability field for both peridotite and pyroxenite, separated by an episode of decompression. The second HP event is dated herein at c. 450 Ma. We have not, however, obtained a reliable date for the first HP episode and cannot determine whether or not it occurred during the mid-Proterozoic, as believed by one of us (H.v.R.), or during the Caledonian orogenic cycle.
The melt-extraction event that resulted in the formation of depleted harzburgite and dunite may have been recorded by the Re–Os isotope patterns of Os-bearing sulfides from a single thick section of sample Fr00-19. Sulfides from peridotites normally do not plot as linear arrays on Re–Os isochron diagrams, probably because they commonly contain two or more sulfide generations (Alard et al., 2000; Pearson et al., 2002). This situation may apply to samples Fr-8 and Fr00-8. The roughly linear Re–Os data from sample Fr00-19 and the nearly isochronous data from a single thick section of this rock are unusual, and could date this event at c. 3·0 Ga. A similar Archean depletion event in WGR peridotite has been proposed by Brueckner et al. (2002). Despite these suggestive data, the linear array could simply be a mixing line, and so the timing of the depletion event must be regarded as provisional even though it must have occurred at some time during the evolution of the sub-Baltic mantle.
The mid-Proterozoic (TDM) Sm–Nd model ages from clinopyroxene suggest that the mantle was fertilized by the intrusion of pyroxenite-related melts probably causing metasomatism of adjacent peridotite. Although the trace element patterns of clinopyroxene could have been produced by several metasomatic agents, including fluids derived from the adjacent crustal rocks, the similarity of the trace element patterns in the peridotite to those of nearby pyroxenite suggests that the fertilization event was the result of the intrusion of silicate melts. If the melts are assumed to have originated from the depleted mantle, the event is dated at mid-Proterozoic by the TDM Sm–Nd model ages of 1·4–1·8 Ga from most Friningen clinopyroxene and whole-rock analyses. This event could have coincided with a synchronous event in the WGR, suggesting a similar mantle history for both terranes. However, unlike the WGR, where metasomatism was accompanied, or followed shortly thereafter, by mid-Proterozoic metamorphism to form c. 1·6 Ga primary (M1) garnet-bearing assemblages (Jamtveit et al., 1991; Brueckner et al., 2002), the Friningen peridotite may not have recrystallized to garnet-bearing assemblages until the Caledonian Orogenic Cycle. This conclusion is based on a lack of a Proterozoic Sm–Nd mineral age from the Friningen peridotite, although future work may yet yield such an age from garnet grains derived from an M1 assemblage. Fertilization appears to have affected a relatively limited portion of the sub-Baltic Shield mantle, judging by the relatively small numbers of garnet- and/or clinopyroxene-bearing peridotite bodies described in the HP nappes of the SNC. However, a systematic search for these assemblages has not been undertaken and it should be remembered that most peridotite bodies in the WGR are also highly depleted dunite and harzburgite.
The events described above occurred in the mantle and are assumed to be linked to the evolution of the overlying Baltic crust. The metamorphism 450 Myr ago that produced M2 garnet-bearing assemblages within the peridotite, as well as within the host country rocks, are believed to mark the time when fragments of this mantle were introduced as discrete bodies into a slice of the Baltic continental crust as that crust was subducted deeply into the mantle. This model would explain the presence of prograde textures within both the peridotite and country rock eclogite-facies assemblages (Van Roermund, 1985, 1989) as well as provide a plausible model for the presence of peridotite within the SNC (Brueckner & Medaris, 1999). The collision thought to be responsible for this subduction was originally assumed to be the Finnmarkian Orogeny, a collision between the western edge of Baltica and an island arc [the Virisen Arc of Stephens & Gee (1989)] during which Baltica underthrust the arc. However, the Finnmarkian Orogeny occurred at c. 500 Ma. Furthermore, the assumed arc–continent geometry would have resulted in a mantle wedge above subducted Baltica composed of oceanic lithosphere that subsequently was contaminated by subduction-related fluids. The results of this study suggest otherwise. The Friningen garnet peridotite body appears to be fertilized mantle that was contaminated in the Proterozoic, too long ago for it to be the suboceanic mantle of either Iapetus or the Aegir ocean.
The apparent contradictions between the models and the data presented in this paper can be resolved if the c. 450 Ma collision involved Baltica and a rifted fragment of Baltica, i.e. a microcontinent. The microcontinent presumably was underlain by subcontinental mantle that rifted away from Baltica along with the overlying crust. Fragments of this subcontinental mantle could then be incorporated into the subducted margin of Baltica as it underthrust the microcontinent. This model can accommodate an earlier collision between this rifted fragment of Baltica and an island arc (i.e. the Virisen Arc), an event that could have been the 500 Ma Finnmarkian Orogeny. The Finnmarkian also involved deep crustal subduction, the formation of eclogite, and the intrusion of peridotite (Mørk et al., 1988; Kullerud et al., 1990). Fifty million years later the resultant composite terrane could have collided with western Baltica during this as yet unnamed orogeny, which we propose be called the Jämtlandian Orogeny (see also Brueckner & Van Roermund, 2001, 2002, 2003; Van Roermund & Brueckner, in prep.).
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APPENDIX: METHODS AND TECHNIQUES |
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Electron microscopy was performed at EMSA (the Utrecht University Centre for Electron Microscopy and Structure Analysis). Electron microprobe (EMP) analyses were conducted at the microprobe laboratory of the Institute of Earth Sciences using a Jeol JXA-8600 Superprobe, both at Utrecht University (Table 1). Further EMP analyses on different samples were conducted using a CAMECA SX50 at GEMOC, Macquarie University. These minerals were analyzed for selected trace elements by laser ablation microprobe–inductively coupled plasma mass spectrometry (LAM–ICPMS) using a Merchantek LUV 266 LAM attached to an Agilent 7500 ICPMS system at GEMOC, Macquarie University (Table 2). The external standard was the NIST 610 glass, and Ca was used as the internal standard. The technique has been discussed in detail by Norman et al. (1996
, 1998).
Some samples were crushed and pure mineral separates obtained through magnetic separation and hand picking techniques. The separates were leached in hot, concentrated HNO3 and HCl solutions and a cold, dilute HF solution. The samples were then spiked, dissolved, and Sr, Nd and Sm were separated using standard laboratory techniques. Isotopes were analyzed by thermal ionization mass spectrometry (TIMS) using a VG 54-30 system at Lamont–Doherty Earth Observatory of Columbia University and on an NU Plasma ICPMS system at GEMOC, Macquarie University (Table 3). Sm–Nd isochrons were calculated using the Isoplot/Ex program of Ludwig (1998). 147Sm/144Nd errors are assumed to be 0·3%. The errors in 143Nd/144Nd were estimated using the relationship (X2 + Y2)1/2 where X is the standard deviation (2) of 21 replicate standard analyses (0·000028) of the La Jolla Standard obtained at the same time as most of the data were measured and Y is the standard error (2) of each individual run.
Sulfides were analyzed for Re–Os using a Merchantek LUV266 LAM attached to an NU Plasma multicollector ICPMS system at GEMOC, Macquarie University, Australia (Table 4). The analytical technique, and its precision and accuracy, have been described in detail by Pearson et al. (2002). Os and Pt abundances were estimated semiquantitatively by comparison with the synthetic PGE-A sulfide standard that is used to calibrate the Os isotope analyses; Re/Os was measured directly.
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ACKNOWLEDGEMENTS |
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We thank D. A. (Tony) Carswell for sharing his remarkable knowledge of eclogite and peridotite during our field trip to the SNC in 2000. We also thank W. L. Griffin, L. Gordon Medaris, Jr and Gareth Davies for critical reviews that enabled us to improve this paper significantly. This work was funded by grants from the Research Foundation of the City University of New York (Nos 61256 and 68229) and grant EAR 00-00937 from the National Science Foundation to H.K.B. and Charlie Langmuir. H.K.B. would like to thank the faculty and staff of Macquarie University for their support and help and the use of their equipment during his sabbatical there in 2000. This paper is contribution 320 from the ARC National Key Centre for Geochemical Evolution and Metallogeny of Continents (GEMOC) and contribution 6472 from Lamont–Doherty Earth Observatory of Columbia University.
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FOOTNOTES |
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* Corresponding author. E-mail: hermanvr{at}geo.uu.nl.
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