Timing and characterization of recurrent pre-Sveconorwegian metamorphism and deformation in the Varberg–Halmstad region of SW Sweden

Precambrian Research 98 (1999) 173–195 www.elsevier.com/locate/precamres

Timing and characterization of recurrent pre-Sveconorwegian metamorphism and deformation in the Varberg–Halmstad region of SW Sweden ˚ ha¨ll b Catherine A. Christoffel a, James N. Connelly a, *, Karl-Inge A a Department of Geological Sciences, The University of Texas at Austin, Austin, TX 78712, USA b Earth Sciences Centre, Go¨teborg University, Box 460, SE-405 30 Go¨teborg, Sweden Received 14 October 1998; accepted for publication 27 May 1999

Abstract ¨ tran terrane of southwestern Sweden is a consequence of westward crustal growth away from an Archean core The A and records both Gothian (c. 1.75–1.55 Ga) and c. 1.0 Ga Sveconorwegian (Grenvillian) tectonism. In spite of widespread consensus on the first-order aspects of this westward growth model, there is a current debate as to whether the bulk of ¨ tran terrane should be attributed to Gothian or Sveconorwegian events. the deformation and migmatization in the A U–Pb zircon, titanite, monazite and rutile ages from gneiss samples and three tectonostratigraphic groups of dykes in the Varberg–Halmstad region constrain the timing of major tectonothermal events. A felsic orthogneiss yields a protolith age of 1664±7 Ma. Mafic dykes of a first dyke group (Steninge dykes) intruded after the first recognized deformation (D1) and were subsequently metamorphosed at 1654±9 Ma. The Steninge dykes constrain D1 to the 1664–1654 Ma interval and D2 at 1654±9 Ma, requiring the D1/D2 gneissosity to have formed during Gothian orogenesis. Steninge dykes and migmatitic D1/D2 structures in the host gneisses are cross-cut by less deformed dykes of a second dyke group, one regional granitic suite dated at 1426+9/−4 Ma and a pegmatite swarm at 1399+7/−6 Ma. The second group of dykes therefore confirms a pre-1.43 Ga age for the main D1/D2 gneiss-forming event(s) consistent with previous models of regional Gothian deformation in this terrane. Group 2 magmatism also includes the Varberg Charnockite–Granite Association, from which a sample of charnockite is now dated at 1399+12/−10 Ma. In addition, metamorphism at 1438+12/−8 Ma resulted in new zircon growth in a mafic gneiss. Together, the Group 2 magmatic rocks and metamorphic zircons represent an interorogenic, 1.44–1.38 Ga thermo-magmatic episode in the Varberg–Halmstad region that has c. 1.4 Ga equivalents elsewhere in southern Sweden. Previous models of high heat flow and comprehensive metamorphism during a thermomagmatic Hallandian event at c. 1.4 Ga is thus supported. Sveconorwegian deformation and recrystallization (D3–D4) occurred before 946+6/−4 Ma, as constrained by a post-kinematic third group of dykes. U–Pb ages for monazite (948±9 Ma), titanite (935±7 Ma and 932±5 Ma) and rutile (878±9 Ma) reflect Sveconorwegian cooling after peak metamorphism. During this stage, U–Pb ages and closure temperatures suggest slow cooling rates of 5–11°C/m.y. from 948 to 932 Ma and 2.5–5°C/m.y. from 932 to 878 Ma, attributed to late-stage erosion and isostatic uplift. The lack of regional penetrative migmitization during the Sveconorwegian granulite conditions is consistent with metamorphism of previously dehydrated rocks (during Gothian ¨ tran terrane. © 1999 and Hallandian times) during eastward Sveconorwegian thrusting of western segments over the A Elsevier Science B.V. All rights reserved. Keywords: Gothian orogen; Mesoproterozoic evolution; Southwest Scandinavian Domain; Southwest Sweden; Sveconorwegian orogen; U–Pb geochronology

* Corresponding author. Present address: Conoco, 10 Desta Drive Suite 100W, Midland, TX 79705, USA. Fax: +1-512-471-9425. E-mail address: [email protected] (J.N. Connelly) 0301-9268/99/$ – see front matter © 1999 Elsevier Science B.V. All rights reserved. PII: S0 3 0 1- 9 2 68 ( 9 9 ) 0 00 4 6 -7

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1. Introduction The assembly of continents in the North Atlantic region during the Paleoproterozoic resulted in a Laurentia–Baltica supercontinent with a presumed contiguous margin, segments of which are presently preserved in eastern Laurentia and western Baltic. Post-assembly consumption of oceanic crust along this continental margin between 1.75 and 1.55 Ga fuelled intermittent magmatism and arc accretion that created most of the lithotectonic elements of SW Sweden and adjoining part of SE Norway. This 200 m.y. period of westward growth in Baltica (present-day co-ordinates) has been termed the Gothian orogeny and is roughly equivalent in time to the Labradorian ˚ ha¨ll and Gower, orogeny of northeast Laurentia (A 1997). The Gothian growth phase ended in SW Sweden by 1.55 Ga, as defined by late- to posttectonic dykes emplaced across existing gneissic ˚ ha¨ll, 1996). fabrics (Connelly and A These subduction-related processes were followed by a protracted period of interorogenic magmatism that occurred intermittently throughout SW Sweden between 1.50 and 1.20 Ga. Subsequent Sveconorwegian orogenesis (1.15– 0.90 Ga), related to Grenvillian continent–continent collision, caused variable reworking of existing rocks but added only minor new crustal components in the eastern Swedish segment of the Sveconorwegian orogen. Given the extensive spatial overlap of the Gothian and Sveconorwegian orogens, the 1.50–1.20 Ga interorogenic intrusive rocks serve as useful structural markers to distinguish between Gothian and Sveconorwegian deformation. Despite recent advances and widespread acceptance of Gothian magmatism, accretion and orogenesis, debate continues regarding the style and extent of pre-Sveconorwegian deformation in ˚ ha¨ll, 1996; Johansson et al., southern Sweden (A ˚ 1996; Aha¨ll et al., 1997a; Mo¨ller and So¨derlund, 1997) where Sveconorwegian reworking locally reached granulite facies conditions (Johansson et al., 1991; Mo¨ller and So¨derlund, 1997). Discrimination of Gothian and Sveconorwegian events is further complicated by hints of a c. 1.4 Ga metamorphic event in the Varberg area

[‘Hallandian’ of Hubbard (1975)]. This study integrates field work and U–Pb geochronology to evaluate pre-Sveconorwegian magmatism, metamorphism and deformation in the well-exposed Varberg-Halmstad region of SW Sweden, where high-grade Sveconorwegian reworking has masked the earlier history. This advance in our understanding of the pre-Sveconorwegian history will assist in developing more accurate tectonic models for the crustal growth in this terrane and evaluating variations in tectonic style along the Laurentia– Baltica supercontinent margin.

2. Regional geology 2.1. Gothian orogen Eastward subduction beneath the Svecofennian margin of Baltica resulted in a major granitoid belt, the Transscandinavian Igneous Belt ( TIB; Fig. 1; e.g. Gorbatschev and Bogdanova, 1993). The belt was principally formed during 1.81– 1.77 Ga magmatism ( TIB 1) but includes westward-younging magmatic phases that intruded between 1.72 and 1.66 Ga ( TIB 2 and TIB 3; ˚ ha¨ll and Larson, in Larson and Berglund, 1992; A press; So¨derlund et al., 1999). These TIB events overlap temporally with oceanward-stepping, subduction-related Gothian magmatism and arc accretion that occurred between 1.75 and 1.55 Ga and formed westward younging segments from the TIB ˚ ha¨ll to the Oslo Rift ( Fig. 1; Connelly et al., 1996; A and Gower, 1997; Brewer et al., 1998). The lack of evidence for subsequent convergent margin processes in the Gothian orogen, combined with widespread and repeated emplacement of 1.50–1.20 Ga rock associations typical of intracra˚ ha¨ll and Connelly, 1998), tonic magmatism (A imply that the pre-existing early Mesoproterozoic margin of Baltica had stepped westward prior to 1.50 Ga. This is consistent with the presence of pre-1.60 Ga crust in the core of SW Norway (‘Telemark craton’; Ragnhildstveit et al., 1994) and revised ages for two supracrustal units on each side of the Skagerrak/Oslo Rift that collectively favour docking of the ‘Telemark craton’

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rane (Fig. 1 inset) is dominated by TIB granites that remained virtually undeformed until the Sveconorwegian amphibolite-facies overprinting (Larson, 1996; Page et al., 1996b; So¨derlund et al., 1999). To the south, complex, polyphase ortho¨ tran terrane (Fig. 1) were gneisses of the A deformed and metamorphosed during both the Gothian and Sveconorwegian orogenies according to the models of Larson et al. (1986) and Connelly et al. (1996). These two terranes were grouped together as the Eastern Segment until it was recognized that they experienced distinct tectonic histories (Larson, 1996; Larson et al., 1998). To the west, the Idefjorden terrane is interpreted to be a distinct Gothian growth zone assembled by c. ˚ ha¨ll and Gower, 1997; Brewer et al., 1.58 Ga (A 1998). The Mylonite Zone bounds the Idefjorden terrane to the east ( Fig. 1), representing a major Sveconorwegian shear zone (e.g. Park et al., 1991; Page et al., 1996a; Stephens et al., 1996; Berglund, 1997) that likely exploited a pre-existing structure that separated older Gothian rocks in the east ˚ ha¨ll, from younger Gothian rocks to the west (A ˚ 1995; Aha¨ll and Gower, 1997). ¨ tran terrane 2.2. A Fig. 1. Lithotectonic map of southern Sweden showing the Steninge area and the Varberg charnockite–Granite Association ¨ tran terrane is delimited by the Mylonite (MZ) (CGA). The A and Protogine (PZ ) zones. The inset map shows major units in Scandinavia and subdivision of the largely overlapping Gothian/Kongsbergian and Sveconorwegian provinces in southern Norway and SW Sweden (grey ornament). The ¨ tran (A ¨ ) and Klara¨lven ( K ) terranes are Idefjorden (Id), A located east of the Permian Oslo Rift, whereas the Kongsberg ( Ko), Bamble (B), Telemark ( T ) sectors are to the west. The Western Gneiss Region ( WGR) lies north of the Caledonides. TIB denotes the Transscandinavian Igneous Belt.

˚ ha¨ll et al., with Gothian units at c. 1.55 Ga (A 1998, 1999). The Gothian orogen and a previously undeformed segment of western TIB rocks were reworked during the Sveconorwegian Orogeny to form a 250 km wide belt that has been divided into three lithotectonic terranes: the Idefjorden, ¨ tran and Klara¨lven terranes [Fig. 1; as defined A ˚ ha¨ll and Gower (1997)]. The Klara¨lven terby A

¨ tran terrane (Fig. 1) is dominated by The A penetratively migmatized gneisses that experienced amphibolite- to granulite-grade metamorphism during the Sveconorwegian Orogeny. Local preservation of eclogites attests to higher pressures and rapid exhumation (Mo¨ller, 1998). Minor occurrences of well-banded units have been interpreted as being supracrustal in origin (Samuelsson et al., ˚ ha¨ll, 1995). The oldest orthogneisses have 1988; A yielded protolith ages between 1.70 and 1.61 Ga and are penetratively deformed (Larson et al., 1990; Johansson et al., 1993; Connelly et al., 1996; Johansson, 1998). Most orthogneisses are granitic in composition, but intermediate and mafic units commonly occur. In the east, 1.66 Ga gneissic granites have been correlated with the unmigmat¨ tran ized 1.68–1.65 Ga TIB 3 granites east of the A terrane (Connelly et al., 1996; Larson et al., 1990). Late- (1.55 Ga) and post-Gothian (1.45, 1.38 and 1.22 Ga) granitic magmatism were intermittent and widespread throughout this terrane

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(Johansson et al., 1993; Berglund and Connelly, 1994; Welin, 1994; Connelly et al., 1996; Lindh, ˚ ha¨ll et al., 1997b; Andersson, 1997). Mafic 1996; A dykes have been dated at c. 1.57 (Ask, 1996), 1.41 (Lundqvist, 1996) and 1.2 Ga (Sm–Nd; Johansson and Johansson, 1990), but older mafic dykes are present (Steninge dykes, see below). East-vergent, Sveconorwegian thrusting related to regional contraction had ceased by 984 Ma in the entire orogen (Romer and Smeds, 1996; cf. ¨ tran Berthelsen, 1980). Emerging data for the A terrane imply that stacking of thrusted crustal slices had induced peak metamorphic conditions by 969±14 Ma, as recorded by new zircon growth (Johansson et al., 1998; Cornell et al., 1996; So¨derlund, 1996; Wang et al., 1998). Approximately 960 Ma Ar–Ar hornblende ages (Page et al., 1996a; Wang et al., 1996) and a 956 Ma Pb–Pb zircon evaporation age for a postmigmatitic granite dyke (Mo¨ller and So¨derlund, 1997) are consistent with this scenario. Cooling during exhumation of mid-crustal rocks is compatible with 945–925 Ma U–Pb titanite (Connelly et al., 1996; Wang et al., 1998) and Ar–Ar hornblende ages (Page et al., 1996a; Wang et al., 1996). Some of these ages and c. 920 Ma U–Pb titanite ages in the Mylonite Zone ( Fig. 1; Johansson and Johansson, 1993) have been related to extension that apparently exploited earlier contractional shear zones. Further cooling due to late-stage exhumation is recorded by c. 905 Ma Ar–Ar muscovite ages (Page et al., 1996a). A number of recent studies have attempted to partition structural and metamorphic elements between Sveconorwegian and pre-Sveconorwegian ˚ ha¨ll, 1995; Connelly et al., 1996; Lindh, events (A ˚ ha¨ll et al., 1997b; 1996; So¨derlund, 1996; A Andersson, 1997; Mo¨ller and So¨derlund, 1997; Johansson, 1998; Wang et al., 1998). In the north, a 1.61 Ga aplitic dyke cross-cuts 1.66 Ga orthogneisses and therefore requires that at least one component of migmatization and deformation ¨ tran terrane was Gothian within this part of the A (Connelly et al., 1996). Ancillary evidence for Gothian deformation comes from a 1.47 Ga aplite that cuts migmatitic fabrics, and includes a misoriented xenolith of host orthogneiss (Connelly et al., 1996). Disputes still exist, however, regarding the

timing of the migmatization and regional gneiss¨ tran terrane formation in the southern part of the A ˚ ha¨ll, 1996; Johansson et al., 1996; A ˚ ha¨ll et al., (A 1997a; Mo¨ller and So¨derlund, 1997) where granulite facies assemblages are partly preserved (e.g. Johansson et al., 1991; Mo¨ller, 1998). For example, field observations and U–Pb (So¨derlund, 1996), Sm–Nd (Johansson et al., 1991; Johansson and Kullerud, 1993) and Ar–Ar ages ( Wang et al., 1996) have been interpreted to date a dominantly Sveconorwegian gneiss-forming event. In contrast, detailed field mapping combined with U–Pb geochronology by Larson et al. (1990), Berglund and ˚ ha¨ll (1995), Connelly et al. Connelly (1994), A ˚ ha¨ll et al. (1997b) and this study, are (1996), A taken as evidence that migmatization and associated deformation occurred mainly before the Sveconorwegian, primarily during Gothian orogenesis. Exceptions exist in Sveconorwegian shear zones where the pre-existing gneissosity has been obliterated or variably obscured by rotation and/or recrystallization. Despite the different schools of thought regarding the timing of the major gneiss-forming deformation, it is widely accepted that granulite-facies conditions were attained in southern parts of the ¨ tran terrane during the Sveconorwegian Orogeny A (e.g. Johansson et al., 1991; Mo¨ller and So¨derlund, 1997; Mo¨ller 1998). To complicate matters further, preservation of charnockites outside Sveconorwegian shear zones in the Varberg area ˚ ha¨ll et al. (1997b) to deduce that highprompted A grade metamorphism occurred at c. 1.4 Ga, returning full circle to Hubbard’s interpretation (Hubbard, 1975) that these charnockites formed at c. 1.4 Ga, attendant with the emplacement of the Varberg Charnockite–Granite Association. Recent data indicate that c. 1.4 Ga ’Hallandian’ charnockitization also occurred in the Halmstad area (Johansson, 1998).

3. Evolution of the Varberg–Halmstad region This study focuses on the Varberg Charnockite– Granite Association (CGA; Hubbard, 1975; Talbot ˚ ha¨ll et al., 1997b) and the and Heeroma, 1989; A well-exposed Steninge area between Varberg and Halmstad ( Fig. 2, inset), an area previously inves-

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Fig. 2. Geological map of the Steninge area modified from Caldenius et al. (1966) with the location of geochronological samples indicated with solid triangles.

tigated by Quensel (1951), Caldenius et al. (1966) and So¨derlund (1996). This coastal area comprises early gneisses, foliated/lineated granites, at least three tectonostratigraphic groups of dykes and four phases of ductile deformation (D1–D4; Table 1). The Group 2 dykes are of particular interest as they serve as structural markers between early gneiss-forming events and later deformation that includes the high-grade Sveconorwegian metamorphism. 3.1. Lithologies and deformation The oldest rocks comprise granitoid orthogneisses and well-banded, felsic gneisses of possible

supracrustal origin. Mafic gneisses are interpreted as largely coeval but may represent younger rock units. Early deformation (D1) imparted the first gneissic fabric with feldspar-rich leucosomes. A swarm of mafic dykes, up to a metre wide, form the first dyke suite (Steninge dykes; Table 1). These abundant, regionally recognized dykes are parallel to sub-parallel to the D1 fabric but do not show any evidence of D1 deformation or related migmatization. They were affected by D2 deformation during at least amphibolite-grade conditions that produced their first fabric and metamorphic zircon (described and dated below). D2 may also have caused tight to isoclinal folding, but this folding cannot unequivocally be assigned to this deforma-

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Table 1 Emplacement and deformation history in the Varberg–Halmstad area, SW Swedena Event/rock unit

Age (Ma)

Deposition; mafic and granitoid intrusion Supracrustal gneiss at Stensjo¨ Sa¨rdal orthogneiss (Sa¨rdal locality) Gneiss at So¨ndrum (Johansson, 1998) D1 Gothian: probably amphibolite facies Mafic dyke intrusions (Group 1 dykes, see text) Steninge dykes (Sa¨rdal locality; metamorphic zircons date D2 recryst.) D2 Gothian; amphibolite facies Resulted in a penetrative planar gneissic D1/D2 fabric Granite intrusion Hinneryd granite (Lindh, 1996) Skipa˚s granite Hallandian episode (incl. Group 2 dykes, see text) Zircon growth in mafic gneiss (Ga˚sanabbe locality) Irregular granitic dykes/neosomes (Sa¨rdal locality) Stensjo¨ pegmatites (1409–1399 Ma); this study Varberg Charnockite–Granite Association: Varberg granite ˚ ha¨ll et al., 1997b) Torpa granite (A Tja¨rnesjo¨ granite (Andersson, 1997) c. 1370 Mafic dyke intrusions D3 Sveconorwegian; high-grade (Johansson et al., 1998) Strong to moderate, linear fabric; penetrative recrystallization Only local formation of migmatitic leucosome D4 mainly chevron folding, subhorizontal axial planes Post-kinematic dyke (Group 3 dykes, see text): Pegmatite (at Stensjo¨) Mafic dykes

– 1664±7 1661±13

1654±9

1548±10 – 1438+12/−8 1426+9/−4 1399+7/−6 1399+7/−6 1380±6 – ≥969±14

946+6/−4 –

a U–Pb ages from this study constrain the timing of events. All ages are U–Pb zircon ages. Rocks dated in this study are in bold text.

tion. Together, the D1 and D2 events resulted in a composite planar foliation in the gneisses ( Table 1) that predates Group 2 dykes. A suite of granitic intrusions that typically exhibit nebulus margins and are largely unmigmatized occur throughout the region. Although this suite contains a marked linear fabric, their relationship with gneissic fabrics of the host gneisses suggests emplacement after the first phase of migmatization such that they represent a separate intrusive event from the orthogneiss protoliths. The lineated Skipa˚s granite found in the Steninge area ( Fig. 2) is a member of this intrusive phase, but dating was hampered by complex zircon systematics (see below). It is not clear whether it preor postdates the D2 event ( Table 1). We tentatively correlate this unit with the Hinneryd granite

(1548 Ma; Lindh, 1996), consistent with observations that it appears to share the first fabric with the Group 2 rocks ( Table 1). Group 2 dykes include two types of granitic rocks and a suite of mafic dykes. Many of these dykes are subparallel to the D1/D2 gneissic fabric, but serve as useful structural markers between D1/D2 and subsequent deformation events where they were emplaced at high angles to the early fabric. One type of granitoid dyke has irregular margins and grades from coarse-grained, layerparallel neosomes in the orthogneiss to metre-wide bodies typically emplaced discordant to the D1/D2 gneissosity. A sample of this type was collected where it unequivocally cuts the D1/D2 gneissosity and has been dated. The other felsic member of the Group 2 dykes occurs as straight-walled peg-

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Fig. 3. High-angle contact between planar D1/D2 gneissosity including migmatitic leusosomes and a Group 2 pegmatite dyke. This metre-wide dyke contains a marked sub-horizontal linear component such that the vertical face exposed in this picture appears weakly deformed. A rock hammer is shown for scale.

matites that cut the D1/D2 gneissosity (Fig. 3); one sample of this type has also been dated. Mafic members of the Group 2 dykes contained no minerals suitable for U–Pb geochronology. D3 represents the first regional deformation event to have affected the Group 2 dykes. It produced predominantly linear fabrics defined by mineral alignment that coincides with F3 fold hinges, but no migmatitic leucosomes occur in the Group 2 dykes or along their margins. The D3 lineation is most easily distinguished in the Group 2 dykes where they are discordant to the earlier D1/D2 gneissic fabrics in the migmatitic orthogneisses. Abundant pinch and swell structures are also outlined well by the generally straight-walled Group 2 pegmatites (cf. So¨derlund, 1996). D3 locally produced open to tight F3 folds with an axial planar foliation ( Table 1). Planar D3 structures are typically subparallel to D1/D2 planar fabrics and can therefore only be reliably distinguished in areas of greater obliquity. D3 was followed by penetrative recrystallization resulting in granoblastic textures in most existing rock types and abundant post-kinematic garnet blastesis in rocks of suitable composition. Locally preserved eclogites attest to high-pressure conditions, a clock-wise P–T–t history and rapid granulite-facies decompression after 969±14 Ma metamorphism (Johansson et al., 1991, 1998; Mo¨ller and So¨derlund, 1997; Mo¨ller, 1998). D3 is

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locally overprinted by D4 deformation, resulting in reorientation of D1–D3 fabrics about open and chevron-type F4 folds with subhorizontal axial planes. Group 3 dykes comprise post-kinematic pegmatitic and mafic dykes that cross-cut all earlier fabrics. Waning static metamorphism and minor brittle deformation affected some of these dykes and suggest a late-Sveconorwegian emplacement. An undeformed pegmatite lens in a deformed mafic dyke was sampled to constrain the final phase of deformation in the area. In the northern part of the study area, the Varberg CGA occupies the same tectonostratigraphic position as the Group 2 dykes ( Table 1). The marked contrast between the moderate fabric in this suite and the gneissic layering in the surrounding gneisses is interpreted to imply that emplacement of this intrusive suite post-dates the gneiss-forming event in the area (cf. Hubbard, ˚ ha¨ll et al., 1997b). Migmatitic leucosomes 1975; A occur but are mainly confined to abundant shear zones in the northern part of the Varberg CGA, where branches of the Mylonite Zone developed along and near the margin of this intrusive complex ˚ ha¨ll, 1995). (e.g. Talbot and Heeroma, 1989; A Given the importance of this intrusion in constraining the major gneiss forming event and timing of the charnockite formation, a sample of charnockite was dated. 3.2. U–Pb geochronology Having established the relative timing of events through field relationships ( Table 1), U–Pb geochronology was employed to constrain: (1) the protolith age of the orthogneiss, (2) the ages of the first, second and third groups of dykes, (3) the timing of major deformation events, and (4) the timing of the regional gneiss-forming event. We also determined a Sm–Nd mineral isochron age for a sample of mafic gneiss to compare this age with U–Pb ages. All zircon and titanite fractions were carefully hand-picked from mineral concentrates using a binocular microscope. Zircon fractions were further scrutinzied by a petrographic microscope (condenser lens inserted ), characterized by cathodoluminescence, extensively

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Table 2 U–Pb isotopic dataa Fraction

Weight (mg)

Concentration

Measured

U (ppm)

Common PbT (pg)

206Pb/ 204Pb

3 3 6 12 124 213 6 156 189

3828 6094 17 796 6274 1064 716 1595 588 622

0.1737 0.1563 0.1283 0.1264 0.1222 0.1244 0.1221 0.8156 0.7120

0.29543 0.29450 0.27954 0.27972 0.27785 0.28076 0.27747 0.15570 0.15501

262 100 430 86 98 112 124 58 56

4.1526 4.1488 3.8627 3.8565 3.8357 3.8640 3.7561 1.5091 1.5005

366 132 590 118 152 168 188 78 64

0.10194 0.10217 0.10022 0.09999 0.10012 0.09982 0.09818 0.07029 0.07020

4 32 7 8 8

14 350 4326 11 545 18 454 13 291

0.1565 0.2188 0.1425 0.1354 0.1129

0.27782 0.27608 0.25766 0.24998 0.23973

102 96 128 114 84

3.8185 3.8410 3.4474 3.2828 3.0963

144 144 138 132 104

6 5 3 8 23 26 16 8 10 41 62

32 692 22 541 42 962 46 575 2031 4281 14 071 6228 6564 229 146

0.1798 0.1859 0.1692 0.1691 63.6125 43.1061 16.5857 22.1250 28.0825 0.0309 0.0444

0.24607 0.24326 0.24252 0.24239 0.24310 0.18762 0.15884 0.15895 0.15819 0.14602 0.14944

86 78 72 82 76 54 52 68 40 246 78

3.0510 3.0091 2.9972 2.9999 2.9841 2.0437 1.5499 1.5523 1.5406 1.3758 1.4421

153 26 7

4548 7995 7387

0.0464 0.0456 0.0424

0.22033 0.23866 0.23796

54 56 82

5 125 7 13 136 131

7773 1256 5576 1570 2595 372

0.0444 0.0693 0.0488 0.0570 0.0960 0.1540

0.24114 0.24049 0.23691 0.23094 0.15540 0.15594

7 4 93

3947 16 214 3920

0.0738 0.0625 0.0946

5 114 19 17

5085 1099 5165 7354

4 6 3 2

27 541 13 146 28 338 41 897

PbR (ppm)

Paleosome (sample 1) Z1 1 lg clr prsm 0.004 123 39.7 Z2 1 lg clr prsm 0.006 162 51.6 Z3 med clr el 0.033 192 56.9 Z4 med clr prsm 0.040 111 32.7 Z5 med clr rnd 0.033 226 66.1 Z6 med clr el need 0.043 195 57.5 Z7 1 med clr el euh 0.002 230 67.0 T1 med clr brn frag 0.081 112 28.4 T2 med clr yl 0.104 113 27.0 Mafic dyke, first dyke suite (sample 2) Z1 16 best sm clr 0.012 305 91.7 Z2 sm lt yw ir eq 0.015 512 160.6 Z3 sm clr rnd ir 0.013 409 112.5 Z4 sm clr-frac rnd 0.017 566 149.8 Z5 sm clr incl rnd 0.010 673 167.5 Pegmatite dyke, second dyke suite (sample 3) Z1 lg clr euh frag 0.047 256 68.9 Z2 lg clr euh prsm 0.027 260 69.4 Z3 lg clr euh frag 0.041 231 60.8 Z4 lg clr-frac euh 0.101 240 63.1 M1 1 v lg clr yw blocky 0.014 206 2805.8 M2 3 med clr lt yw rnd 0.028 339 2440.8 M3 1 v lg clr yw blocky 0.022 1065 2589.3 M4 1 v lg clr yw blocky 0.008 578 1846.6 M5 1 v lg clr yw blocky 0.013 524 2095.9 R1 lg clr gold el 0.200 5 0.6 R2 med clr gold el 0.205 4 0.6 Pegmatite dyke, second dyke suite (sample 4) Z3 lg clr rnd eq 0.056 897 192.3 Z1 lg rnd frac 0.012 1152 267.9 Z2 1 clr rnd 0.003 1134 262.4 Mafic gneiss (sample 5) Z1 v sm clr eq rnd 0.007 372 87.4 Z2 sm clr eq rnd 0.031 327 78.4 Z3 v sm clr rnd eq 0.012 212 49.1 Z4 sm clr rnd eq 0.014 97 22.0 T1 lg clr dk brn frag 0.131 275 42.9 T2 lg clr lt yw frag 0.112 43 7.0 Pegmatite dyke, third dyke suite (sample 6) Z1 1 best clr frag 0.010 258 39.9 Z2 1 clr frag 0.025 293 44.6 Z3 1 v lg pk frag 0.181 208 32.1 Skipa˚s granite (sample 7) Z1 lg incl clr euh el 0.005 324 96.5 Z2 med clr rnd 0.019 362 106.5 Z3 med clr incl 0.013 439 120.2 Z4 sm el incl frac 0.022 349 94.9 Varberg Charnockite (sample 8) Z1 HF lg lt brn frag 0.131 59 14.7 Z3 lg clr ang frag 0.113 53 13.2 Z4 1° mag clr frag 0.094 63 15.6 Z5 lg clr ang frag 0.120 64 15.1

Corrected atomic ratiosb 208Pb/ 206Pb

206Pb/ 238Pb

Ages (Ma)

207Pb/ 238U

207Pb/ 235U

206Pb/ 206Pb

207Pb/ 238U

207Pb/ 206Pb

20 18 18 12 12 24 32 24 16

1669 1664 1589 1590 1581 1595 1579 933 929

1665 1664 1606 1605 1600 1606 1583 934 931

1660 1664 1628 1624 1626 1621 1590 937 934

0.09968 0.10091 0.09704 0.09524 0.09367

14 12 30 24 14

1580 1572 1478 1438 1385

1597 1601 1515 1477 1432

1618 1641 1568 1533 1502

112 102 98 110 88 58 52 58 38 242 144

0.08992 0.08972 0.08963 0.08976 0.08903 0.07900 0.07077 0.07083 0.07063 0.06833 0.06999

10 10 10 8 14 10 6 16 10 58 60

1418 1404 1400 1399 1403 1108 950 951 947 879 898

1420 1410 1407 1408 1404 1130 950 951 947 879 907

1424 1419 1418 1420 1405 1172 951 953 947 879 928

2.5883 2.9007 2.8962

66 72 94

0.08520 0.08815 0.08827

6 6 16

1284 1380 1376

1297 1382 1381

1320 1386 1388

104 72 86 102 60 82

2.9757 2.9608 2.9096 2.8090 1.5042 1.5099

122 100 94 110 62 100

0.08950 0.08929 0.08907 0.08822 0.07021 0.07023

20 14 18 24 10 38

1393 1389 1371 1339 931 934

1401 1398 1384 1358 932 934

1415 1410 1406 1387 934 935

0.15751 0.15628 0.15413

68 56 46

1.5345 1.5206 1.5024

64 60 50

0.07066 0.07057 0.07069

18 20 10

943 936 924

944 939 931

947 945 949

0.1287 0.1204 0.0949 0.0952

0.28096 0.27978 0.26666 0.26474

102 126 80 92

3.9406 3.8988 3.6390 3.5685

126 194 120 130

0.10172 0.10107 0.09898 0.09776

22 34 10 14

1596 1590 1524 1514

1622 1613 1558 1543

1656 1644 1605 1582

0.2163 0.2194 0.2537 0.2224

0.22391 0.22023 0.21614 0.21177

46 50 48 40

2.6513 2.5916 2.5186 2.4426

56 62 58 50

0.08588 0.08535 0.08451 0.08366

6 8 6 6

1303 1283 1261 1238

1315 1298 1277 1255

1335 1324 1304 1285

a Codes in fraction numbers are: M, monazite; R, rutile; T, titanite; Z, zircon. b Ratios corrected for fractionation (0.1%/amu and 0.15%/amu for Pb run by static Faraday and peak-jumping ion counting, respectively; 0.07%/amu and 0.15/amu for U run by static Faraday and peak-jumping ion counting, respectively), 1 and 2 pg of laboratory Pb blank for zircon and titanite chemistry, respectively, initial common Pb calculated using Pb isotopic compositions of Stacey and Kramers (1975) and 0.25 pg of U laboratory blank. All fractions were extensively abraded. Two-sigma uncertainties on isotopic ratios, calculated with a modified unpublished error propagation program written by L. Heaman, are reported after the ratios and refer to the final digits. Abbreviations are: ang, angular; brn, brown; clr, clear; dk, dark; el, elongate; eq, equant; euh, euhedral; frac, fractures; frag, fragments; HF, best after HF treatment (see text for discussion); incl, inclusions; ir, irregular; lt, light; lg, large; mag, magnetic tilt on Frantz Isodynamic Magnetic separator; med, medium; need, needles; pk, pink; prsm, prism; rnd, round; sm, small; v, very; yw, yellow.

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abraded and re-evaluated opitically before analyses. These steps ensured that the best-quality grains were used for the purpose of dating specific preSveconorwegian events in a terrane overprinted by Sveconorwegian metamorphism. Appendices A and B outline U–Pb and Sm–Nd analytical procedures, respectively, and sample locations are given below using co-ordinates of the Swedish National Grid. Data are presented in Table 2 with two-sigma errors on all ratios. 3.2.1. Sample 1: Sa¨rdal orthogneiss (629720/130586) A homogeneous paleosome layer in a granitic orthogneiss was collected at Sa¨rdal ( Fig. 2) where abundant leucosomes and veins were present but carefully avoided. This strongly migmatized granitoid rock represents the major rock type in the

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study area and is interpreted as the oldest igneous unit. Static recrystallization has obliterated mineral alignment fabrics, enabling this orthogneiss unit to be extensively quarried as coherent facing stone in spite of the earlier gneissic banding. A mafic Steninge dyke and a discordant Group 2 granitic dyke were also sampled at this locality (see below). Zircon morphologies from the Sa¨rdal orthogneiss sample include elongated and equant grains that range from euhedral to anhedral. Cathodoluminescent (CL) images [Fig. 4(a)–(d )] exhibit clear, straight-sided zoning with only thin rims that might reflect metamorphic overgrowth. The external morphologies and internal zonation are consistent with an igneous origin. Care was taken to avoid grains containing cores commonly observed in the euhedral grains using petrographic and CL microscopy.

Fig. 4. (a)–(d) Images of cathodoluminescence in four zircon grains representative of Sample 1. All grains exhibit straight-edged zonation typical of magmatic growth of zircons in felsic igneous magmas. The appearance of more complex interiors is due to the low angle between the polished surface and internal zonation, in contrast to the high-angles nearer the edges.

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U–Pb isotopic data for five small, clear, colourless fractions ( Z1–Z5) of elongate, euhedral zircons ( Table 2) define a discordia line with intercepts at 1659±5 and 872+71/−66 Ma [probability of fit 12%; Fig. 5(a)]. Single grain fractions Z1 and Z2 plot concordantly at 1664±4 Ma. Considering the overlapping results, we defer to the concordant fractions to define a protolith age of 1664±7 Ma where the error has been increased slightly to include the regressed upper intercept. We interpret this date to reflect the crystallization age of this granitoid orthogneiss, commensurate with the zircons igneous heritage. Data from zircon fractions Z6 and Z7 [Fig. 5(a)] plot above the discordia line and may reflect multi-stage Pb loss and/or zircon overgrowth during one or more metamorphic events. This sample also yielded light orange to brown titanite fragments. Two fractions yield a concordant age of 932±7 Ma [Fig. 5(a), lower inset] that overlaps the lower intercept for the zircon regression (872+71/−66 Ma). 3.2.2. Sample 2: Group 1 (Steninge) mafic dyke (629736/130590) Mafic dykes of Group 1, also referred to here as the Steninge dykes, are abundant in the study area where they are extensively boudinaged. A typical Steninge dyke was sampled at Sa¨rdal (Fig. 2) at the same locality where the host orthogneiss was collected for dating. The dyke is approximately 40 cm wide and traceable for 15 m along a trail of disrupted dyke boudins. The rock is completely recrystallized and has a strong fabric defined by amphibole. The intense recrystallization obscures the relationships between this particular dyke and structures in the gneiss. However, these mafic Group 1 dykes do not have any migmatitic leucosomes that characterize the mafic gneiss in the area (see below) and are part of the suite of mafic dykes that are locally sub-parallel to D1. The sampled dyke is therefore interpreted to postdate an early gneiss-forming event in the region (D1; Table 1). Coarse-grained, granitic interboudin injections are common and similar to granitic Group 2 dykes described below. The sample of Group 1 dykes yielded clear, colourless, spherical to slightly elliptical grains that

are approximately 120 mm in the longest dimension. They exhibit inconsistent irregular internal zoning patterns in CL images, many exhibiting a low luminoscity [Fig. 6(a)–(d )]. These features are indicative of a metamorphic origin for the zircons, an interpretation consistent with the typical lack of primary zircons in mafic dykes. These zircons are also completely dissimilar to zircons recovered from its host rock (the Sa¨rdal gneiss, see above), precluding inheritance from this lithology. A regression through data from four of the five fractions yields an upper intercept of 1654±9 Ma [probability of fit 10%; Fig. 5(b)]. The lower intercept, 941±30 Ma, is interpreted as a product of Sveconorwegian Pb loss. Fraction Z2 plots below the discordia line and is excluded from the regression as it may include an inherited component in this one fraction. Commensurate with the proposed metamorphic origin of these zircons, our preferred interpretation is that the 1654±9 Ma date is interpreted as the age of metamorphic recrystallization of this mafic dyke. This age constrains the lower age for D1 structures and is interpreted to directly date D2, the event that caused metamorphic zircon growth. Given the morphology and internal structure of these zircons, we are confident that they reflect a single population of metamophic origin. However, recognizing the importance of this age to the interpretation below, it is worth reviewing two alternative scenarios. If the zircons are instead of igneous origin, D1 must still pre-date 1654±9 Ma, but the metamorphic assemblage within the dyke could then be significantly younger. An igneous heritage is, however, inconsistent with the external and internal characteristics of this zircon population, which was carefully examined by CL before the best, clear, grains were selected for analyses. A second alternative would have the zircons inherited from the host rock, a scenario in which the date does not define the age of dyke emplacement and thereby constrains neither D1 nor D2. This interpretation is consistent with the overlapping ages of host and dyke, but is inconsistent with the more important observation that the zircons from dyke and host bear no resemblance in form, size or internal structure. Acknowledging the clear metamorphic characteristics of this dyke

C.A. Christoffel et al. / Precambrian Research 98 (1999) 173–195

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Fig. 5. Concordia diagrams with ellipses representing two-sigma errors (Z=zircon; T=titanite; M=monazite; R=rutile). See text for discussion.

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Fig. 6. (a)–(d) Images of cathodoluminescence in four zircon grains representative of Sample 2. All grains imaged exhibit highly irregular, patchy luminescence consistent with a metamorphic origin. These grains are distinct in form and internal characteristics from those grains imaged from Sample 1.

population makes this interpretation even more unlikely as it would require that the sampled mafic dyke inherited only structureless metamorphic zircons from a sea of well-preserved igneous zircons (that we document in the host granitoid orthogneiss above). 3.2.3. Sample 3: Group 2 granite dyke at Sa¨rdal (629734/130590) Coarse-grained granitic veins are common in the migmatized orthogneiss, as described at Sa¨rdal. They vary from localized centimetre-wide lenses to irregular metre-wide dykes that truncate the gneissosity in the host rock. One such dyke was sampled at Sa¨rdal to constrain the early (D1/D2) fabric in the host gneisses that it unequivocally cuts and the pronounced D3 linear fabric that it

contains. Fractions of zircon, monazite and rutile were analysed. Most zircons are large, showing euhedral to rounded prisms, elongate crystals and crystal fragments. CL images indicate primary igneous growth zoning in large, clear, colourless, euhedral, prismatic grains and fragments of grains. Four fractions define a discordia line with intercepts at 1426+9/−4 and 517+307/−227 Ma [probability of fit 26%; Fig. 5(c)]. The upper intercept is interpreted to represent the igneous crystallization of the dyke. The younger lower intercept than those of other samples suggests multi-stage Pb loss. The sample also yielded two morphologies of large, light- to dark-yellow monazite: fragile crystals (with distinct cleavage planes) and blocky, robust grains lacking cleavage planes. Data from

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five monazite fractions define a mixing line with intercepts at 1410±9 and 948±4 Ma [probability of fit 60%; Fig. 5(c)]. Fractions M3–M6 are singlegrain analyses that indicate two separate events. A fraction of dark anhedral monazite (M1) plots nearest the 1410±9 Ma intercept that most likely represents either the closure age of primary magmatic monazite (c. 700°C closure temperature; Heaman and Parrish, 1991) or metamorphic monazite that formed immediately after the crystallization. Either interpretation provides a minimum igneous age that agrees with the 1426+9/−4 Ma zircon crystallization age. Three fractions of robust, blocky monazite crystals provide a concordant Sveconorwegian metamorphic age of 948±4 Ma [M3–M5; Fig. 5(c)]. Multigrain fraction M2 plots discordantly on the line and most likely comprised both monazite generations. Medium to large (75–100 mm), clear, gold to orange, euhedral and multi-faceted rutile grains were rare. A concordant rutile fraction yields a late- to post-Sveconorwegian cooling age of 879±10 Ma [Fig. 5(c)]. The U–Pb system in rutile is easily reset and is expected to reflect either late hydrothermal activity (Machado et al., 1989) or cooling below the closure temperature, estimates of which range from 380 to 420°C depending on the grain radius (Mezger et al., 1989). 3.2.4. Sample 4: Group 2 Stensjo¨ pegmatite (630060/130500) Coarse, K-feldspar-rich pegmatites of the Group 2 dykes are up to 10 m wide and especially abundant in the Stensjo¨ area where well-banded gneisses of presumed metasupracrustal origin occur. Most dykes are approximately 50 cm wide, straight walled and subparallel to the host gneisses, but clearly discordant dykes occur ( Fig. 3). A dyke was sampled at Stensjo¨ (Fig. 2) where it truncates S1/S2 of the host gneiss and has a marked fabric dominated by a linear component (D3). Latestage recrystallization caused local idiomorphic garnets along the margins. Zircon morphologies include large, rounded, clear, colourless, equant grains that are commonly cracked. Analyses of three multigrain zircon fractions yield a discordia line with intercepts at

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1399+7/−6 and 988±33 Ma [probability of fit 12%; Fig. 5(d)]. The 1399+7/−6 Ma age is interpreted to date the crystallization of the Stensjo¨ pegmatite. 3.2.5. Sample 5: Ga˚sanabbe mafic gneiss (630152/130495) A veined and completely recrystallized mafic gneiss occurs in the coastal area where well-banded felsic gneisses of possible supracrustal origin otherwise dominate. A paleosome sample from an approximately 8 m thick mafic unit was collected at Ga˚sanabbe (Fig. 2) for its potential to yield minerals appropriate for U–Pb and Sm–Nd analysis. The rock is a hornblende–clinopyroxene– garnet–biotite–scapolite amphibolite with fabric and leucosomes parallel to the host gneiss. An effort was made to avoid leucosomes during sample collection. Relationships between possible D1 structures in the mafic gneiss and mafic Group 1 dykes have been obscured by recrystallization, but the planar fabric of both the mafic and adjacent gneisses is cut by Group 2 dykes such that this fabric is assigned to D1/D2. This sample yielded rounded, spherical, anhedral grains between 60 and 100 mm in diameter that exhibit irregular zonation and poor luminoscity in CL [Fig. 7(a) and (b)]. These features are typical of metamorphic zircons in mafic rocks. Four multigrain zircon fractions, each with 30– 100 clear, colourless, rounded, equant, anhedral grains between 60 and 100 mm ( Table 2), define a discordia line with intercepts at 1438+12/−8 and 855±100 Ma [probability of fit 22%; Fig. 5(e)]. The upper intercept is interpreted to represent the age of new zircon growth during metamorphism, commensurate with the interpreted origin of these zircons. The lower intercept apparently reflects multistage Pb loss during the Sveconorwegian and more recently. Two fractions of clear, large (>100 mm), light-yellow to dark-brown, anhedral, rounded and fragmented titanite, yield a concordant age of 935±5 Ma [Fig. 5(e)] that constrains the Sveconorwegian cooling history. It is interesting to note that zircons found in this mafic gneiss record an event of zircon growth at 1438 Ma, whereas metamorphic zircons in the younger Group 1 mafic Steninge dykes (see above)

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Fig. 7. (a) and (b) Images of cathodoluminescence in two zircon grains representative of Sample 5. These grains exhibit highly irregular, patchy and typically weak luminescence consistent with a metamorphic origin.

record a metamorphic event that is much older (1654 Ma; Table 1). It appears that conditions were favourable for zircon growth in the Group 1 mafic dyke at the sample site at 1654 Ma (during the D2 event), and, once formed, they remained stable and thus do not record the c. 1.4 Ga event, by either Pb loss or overgrowths. In contrast, there is no evidence for zircon growth in the mafic gneiss at the sampled locality until 1438 Ma, apparently the first time that conditions allowed their formation. That these two mafic rocks record different metamorphic zircon growth histories requires that factors controlling their growth varied locally. If zircon growth in mafic rocks that lack primary zircon is related to the breakdown of Zr-bearing pyroxene, their formation will be controlled by: (1) bulk-rock composition, (2) P–T paths, and/or (3) H O activity. It is possible that the mafic gneiss 2 also contains zircon from the 1654 Ma event and that such grains were successfully avoided by removing zircons that contained cores. 3.2.6. Sample 6: Post-kinematic Group 3 pegmatite (630075/130512) A metre-wide lens-shaped Group 3 pegmatite within a mafic Group 2 dyke was also collected at Stensjo¨ (Fig. 2). The rock is post-kinematic, but minor deformation is locally expressed in this suite as kinked biotites. The pegmatite cross-cuts the fabric in the mafic dyke (S3) and was dated to constrain the timing of Sveconorwegian deformation. Given the post-kinematic emplacement across

the mafic dyke, the pegmatite must have been intruded during or after the last deformation event (D4). Large (1–5 mm) euhedral zircons [Fig. 5(f )] were extracted directly from the outcrop. A single, 4 mm long, light-pink euhedral zircon was crushed, and three inclusion-free fragments were analysed. A concordant fraction [Z1; Fig. 5(f )] provides an igneous crystallization age of 946+6/−4 Ma, and the two near-concordant fractions plot on a reference line to 0 Ma.

3.2.7. Sample 7: Skipa˚s granite (629925/130600) A sample of the Skipa˚s granite was collected to determine a protolith age for the gneissic granite intrusions that are typically less migmatized than the host gneisses (cf. Johansson et al., 1993; Lindh, 1996). Our sample yielded a complex zircon population that included euhedral to subhedral, elongate to spherical colourless zircons. Attempting to isolate primary igneous zircons, four fractions of euhedral to subhedral zircons were analysed but yielded discordant, non-linear data points with 207Pb/206Pb ages between 1.66 and 1.58 Ga [Fig. 5(g); Table 2]. Since we interpret this rock to be younger than D1 (1664–1654 Ma; Table 1) and possibly D2 ( Table 1), we infer that this sample contains inherited zircons from a source rock that is at least 1.66 Ga old. The 1.55 Ga Hinneryd granite (Lindh, 1996 and references therein) that intruded to the southeast may be correlative.

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3.2.8. Sample 8: Varberg Charnockite–Granite Association (633645/128426) The Varberg Charnockite–Granite Association (CGA) (Hubbard, 1975; cf. Quensel, 1951) comprises megacrystic granodiorite-granite ( Torpa granite), medium-grained, equigranular monzonite-granite ( Varberg granite) and minor mixed varieties (cf. Hubbard and Constable, 1980; Talbot and Heeroma, 1989; Johansson and Kullerud, ˚ ha¨ll et al., 1997b). The Varberg granite is 1993; A mainly charnockitic, whereas such orthopyroxene paragenesis occurs only in the interior parts of the Torpa granite. Previous ages include a 1420±52 Ma whole-rock Rb–Sr age for the Varberg CGA ( Welin and Gorbatschev, 1978; recalculated and with two-sigma errors) and a U– ˚ ha¨ll Pb age of 1380±6 Ma for the Torpa granite (A et al., 1997b). Debates over the relationships and subsequent evolution have been complicated by an apparent metamorphic break along the northern ˚ ha¨ll, 1995 and referside of the Varberg CGA (A ences therein) and by different interpretations for the timing of the charnockitization (see Section 4). Homogeneous, charnockitic Varberg granite was collected in the quarry behind the castle in Varberg. This rock has a moderate lineation, and fresh surfaces have a green tinge typical of granulite-facies rocks. The sample yielded large, light beige fragments of euhedral zircons, many containing small but abundant inclusion trails — these could not be avoided in the analyses. Internal zoning, as imaged by CL, is typically absent, although several grains contain a very faint igneous pattern. Three fractions define a discordia line [Z3–Z5; Fig. 5(h)] but are highly discordant and preclude determination of sufficiently precise intercepts. Emersion of two additional fractions (Z1 and Z2) in warm HF for 2 h caused about half the grains to become cloudy. Clear grains were then separated, abraded and analysed normally. These two analyses have the lowest U contents and plot less discordantly than fractions Z3–Z5. Regressed together, the five fractions yield intercepts of 1399+12/−10 Ma and 971±30 Ma [19% probability of fit; Fig. 5(h)], and the upper intercept is interpreted to represent the crystallization age.

187

The HF treatment allowed optical identification of less metamict, low-U grains. These grains underwent less Pb loss during metamorphism, here dated at 971±30 Ma. The lack of distinct igneous zoning in the Varberg granite zircons, a ubiquitous feature of felsic igneous rocks, is attributed to chemical homogenization by thermal diffusion, a process that inherently causes signficant Pb loss. The loss of approximately half the accumulated radiogenic Pb in this zircon population at 971 Ma [Fig. 5(h)] implies that chemical homogenization also occurred at this time. The acceptance that Sveconorwegian metamorphism was responsible for this Pb loss and chemical homogenization does not require that the charnockitization of this unit occurred at this time. It should also be noted that the 1420±52 Ma Rb–Sr whole-rock isochron of the Varberg CGA ( Welin and Gorbatschev, 1978; recalculated and with 2s errors) was not significantly disturbed during the high-grade Sveconorwegian overprinting, in spite of its apparent effects on the U–Pb system of the zircons. The two major constituents in the Varberg CGA are now dated at 1399+12/−10 Ma (mediumgrained Varberg granite) and 1380±6 Ma (meg˚ ha¨ll et al., 1997b). The acrystic Torpa granite; A age difference indicates a prolonged magmatic event as the mixed varieties and gradations between the two rock types point to a simultaneous existence of the two granite magmas (cf. Hubbard, 1975; Talbot and Heeroma, 1989). 3.3. Sm–Nd geochronology Pure titanite, apatite and clinopyroxene were carefully hand-picked from the sample of Ga˚sanabbe mafic gneiss and were analysed for Sm and Nd; garnet was not sufficiently abundant to analyse. The three Sm–Nd analyses define a mineral isochron age of 878±23 Ma (MSWD= 0.23; Table 3; Fig. 8). This result overlaps with previously published Sm–Nd mineral isochrons for the granulite region (0.91–0.88 Ga; Johansson et al., 1991; Johansson and Kullerud, 1993; Wang et al., 1998). Given that these c. 0.90 Ga mineral isochrons are defined by peak metamorphic minerals, it is tempting to interpret these Sm–Nd ages to represent peak

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Fig. 8. Sm–Nd mineral isochron plot for the Ga˚sanabbe mafic gneiss. Error bars represent two-sigma errors. Abbreviations: ap, apatite; cpx, clinopyroxene; ttn, titanite.

metamorphic conditions. However, titanite ages between 950 and 934 Ma in the area (this study), ¨ tran terrane (Connelly et al., and elsewhere in the A 1996; Wang et al., 1998), require that conditions cooled below the closure temperature of Pb in titanite by this time. It is therefore also required that Sm and Nd continued to re-equilibrate both chemically and isotopically after peak conditions. The Sm–Nd mineral ages obtained cannot therefore be used to constrain the age of high-grade conditions (cf. Wang et al., 1998).

4. Discussion 4.1. Gothian events The 1664±7 Ma protolith age for the Sa¨rdal orthogneiss coincides with ages for igneous zircon cores from the same gneiss unit further south

(1661±13 Ma; Johansson, 1998) and with similar ¨ tran terorthogneisses in northern parts of the A rane (1660±5 Ma; Connelly et al., 1996). The 1.66 Ga Sa¨rdal gneiss is also coeval with the 1.66 Ga Horred Formation, an outboard arc ˚ ha¨ll system accreted to Baltica prior to 1.58 Ga (A ˚ ha¨ll, 1996; Brewer et al., 1995; Connelly and A et al., 1998). Still older protolith ages (1.70– 1.69 Ga; Connelly et al., 1996) from gneisses in ¨ tran terrane suggest a northeastern parts of the A stepwise growth in this terrane. ¨ tran orthoThe 1.70–1.69 Ga and 1.66 Ga A gneisses are coeval and have been correlated with the well-preserved TIB 2 and TIB 3 granites that occur immediately east of this terrane (Larson and Berglund, 1992; Connelly et al., 1996). However, a linkage between these alkali-calcic granites in the east and the more juvenile, and possibly ˚ ha¨ll and accreted, crustal units in the west (A Gower, 1997 and references therein) cannot be precluded. Emplacement of the Group 1 mafic Steninge dykes was widespread and is now constrained between 1664±7 Ma (age of the host gneiss at Sa¨rdal ) and 1654±9 Ma (age of metamorphic zircons in the Group 1 dyke sample). The lack of migmatitic veining in Group 1 dykes and their locally oblique relationship with gneissic layering in the host indicate that D1 occurred before the dykes were emplaced. Zircon growth at 1654±9 Ma is interpreted to directly date the metamorphism associated with D2 ( Table 1). This small time period for D1 and D2 permits a continuous D1/D2 event that included injection of the Group 1 mafic Steninge dykes at c. 1.66–1.65 Ga. Given the metamorphic zircon characteristics of the analysed Group 1 dyke (see above), we conclude that the 1654±9 Ma age provides a direct

Table 3 Sm–Nd isotope data for the Ga˚sanabbe mafic gneissa Mineral

Sample size (mg)

Sm (ppm)

Nd (ppm)

147Sm/144Nd

+/−

143Nd/144Nd

+/−

Apatite Clinopyroxene Titanite

6.9 40.3 2.4

39.8 3.4 354.0

240.0 14.1 1546.0

0.10027 0.14596 0.13872

53 114 102

0.511754 0.512018 0.511974

5 6 7

a Nd composition is normalized to 146Nd/144Nd=0.7219. Two-sigma uncertainties on isotopic ratios refer to the final digits.

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U–Pb determination of Gothian metamorphism in ¨ tran terrane (D2). The tectonic cause of this the A event is uncertain, but it may be related to the accretion of the outboard 1.66 Ga Horred arc system onto Baltica (cf. Brewer et al., 1998). The 1426+9/−4 Ma age of the granitic dyke that discordantly cuts the migmatitic Sa¨rdal orthogneiss provides evidence for a pre-1426 Ma migmatization in that the dominant D1/D2 gneissosity in the host gneisses formed prior to the emplacement of these dykes. Independent support for pre-1400 Ma migmatization is provided by the regionally extensive 1399+7/−6 Ma Stensjo¨ pegmatite, that also cuts the dominant D1/D2 gneissosity (Fig. 3). Other members of this swarm have yielded ages of 1409±20 and 1406±6 Ma (So¨derlund, 1996; Johansson, 1998). Thus, amphibolite-facies metamorphism related to D1–D2 deformation must predate 1426 Ma and has been directly dated at 1654±9 Ma in the Steninge area. This evolution is similar to that previously constrained by U–Pb data in northeast¨ tran terrane where 1.61 and ern parts of the A 1.47 Ga dykes are interpreted to provide minimum ages for Gothian deformation in that region (Connelly et al., 1996). These conclusions, based on specific U–Pb geochronology, match the general observation that rocks interpreted as part of the ¨ tran terrane have ‘early gneisses’ throughout the A yielded U–Pb ages within the Gothian interval, i.e. 1.70–1.61 Ga (Johansson et al., 1993; Welin, ˚ ha¨ll et al., 1995; Connelly et al., 1996; 1994; A Johansson, 1998), whereas typically less-migmatized rocks have yielded late-Gothian or still younger U–Pb ages in the 1.55–1.20 Ga interval (Johansson et al., 1993; Berglund and Connelly, ˚ ha¨ll et al., 1997b; 1994; Welin, 1994; Lindh, 1996; A Andersson, 1997). These constraints for regional pre-1.43 Ga migmatization do not support recent proposals that migmatization in these orthogneisses is predominantly Sveconorwegian (cf. Johansson et al., 1996; So¨derlund, 1996; Mo¨ller and So¨derlund, 1997), but are in good agreement with previous models evoking a dominantly Gothian gneiss-forming event (e.g. Larson et al., 1986; Samuelsson et al., 1988; Johansson et al., 1993).

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4.2. Interorogenic ‘Hallandian’ episode Gothian growth in southern Baltica was followed by relatively stable cratonic conditions that were interrupted by the emplacement of 1.50– 1.20 Ga rock suites typical of intra-cratonic mag˚ ha¨ll, 1996; A ˚ ha¨ll and matism (Connelly and A Connelly, 1998). This, and the lack of any evidence for continental margin processes in the Gothian orogen during this interval, collectively imply that the margin of Baltica had moved west of the ˚ ha¨ll et al., Gothian orogen by 1.50 Ga (A ˚ 1998; Aha¨ll and Connelly, 1998). ¨ tran terrane Unique to southern Baltica, the A provides five independent pieces of information for a prolonged, 1.44–1.38 Ga thermo-magmatic episode: (1) growth of metamorphic zircon at 1438+12/−8 Ma (this study), (2) regional formation of 1426+9/−4 Ma irregular granitic dykes (this study), (3) widespread emplacement of pegmatites at 1409–1399 Ma (So¨derlund, 1996; Johansson, 1998; this study), (4) emplacement of the Varberg CGA at 1399–1380 Ma (this ˚ ha¨ll et al., 1997b) and (5) crystallization study; A at c. 1.37 Ga of the Tja¨rnesjo¨ granite east of Varberg (Andersson, 1997). Additional U–Pb constraints come from the 1363±9 Ma old Askim granite in the Idefjorden terrane ( Welin and Samuelsson, 1987), the mafic–felsic 1410±10 Ma ¨ tran Axamo dykes near the eastern margin of the A terrane (Lundqvist, 1996) and a number of 1.45– 1.38 Ga granites in SE Sweden (Johansson et al., 1993 and compilation therein; Kornfa¨lt, 1996). The large scale of the c. 1.4 Ga magmatism in southern Sweden has previously been noted ˚ berg (1988) and Johansson et al. (1993). The by A widespread character of this thermo-magmatic episode is also shown by c. 1.4 Ga isotopic resetting of Rb–Sr whole-rock systems in granites in the Idefjorden terrane ( Welin et al., 1982; Hansen et al., 1989), including the 1503+3/−2 Ma ˚ ha¨ll and Connelly, 1998) Stigfjorden granite (A that yielded a 1416±19 Ma Rb–Sr whole-rock ˚ ha¨ll et al., 1990). isochron age (cf. A The 1.44–1.38 Ga thermo-magmatic episode coincides with the proposed ‘Hallandian Orogeny’ of Hubbard (1975), an event that was widely discounted due to a lack of identifiable deforma-

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tional consequences ( Hubbard, 1976; Samuelsson and Ahlin, 1976; Gorbatschev and Welin, 1980; Hubbard and Constable, 1980; Gaa´l and ˚ ha¨ll et al., 1997b). Granulite Gorbatschev, 1987; A facies in the Varberg area were thought by Hubbard (1975) to be related to emplacement of the Varberg CGA. This field-based interpretation ˚ ha¨ll et al. (1997b), whereas was supported by A Johansson and Kullerud (1993) cited their Sm– Nd mineral isochron ages to assign the formation of charnockites to the Sveconorwegian Orogeny. The temporal mismatch between Sm–Nd mineral ages obtained in the granulite area (910–880 Ma; Johansson et al., 1991; Johansson and Kullerud, 1993; Wang et al., 1998; this study) and the older U–Pb metamorphic ages for peak conditions (c. 969±14 Ma; Johansson et al., 1998) requires that these Sm–Nd mineral isochron ages cannot reflect the timing of granulite conditions (Mo¨ller and So¨derlund, 1997; Wang et al., 1998). Given that recent U–Pb data from the Halmstad area encouraged Johansson (1998) to interpret that charnockitization occurred at c. 1.40 Ga, Hubbard’s model in which charnockites were generated during the ‘Hallandian’ event appears well supported. It is thus concluded that the large-scale, thermomagmatic Hallandian episode at 1.44–1.38 Ga caused scattered but widespread charnockitization of older gneisses in the Varberg–Halmstad region, including crystallization of the partly charnockitic Varberg CGA. Hubbard (1975) and Talbot and Heeroma (1989) have further suggested that the Hallandian episode was associated with deformation, but later studies have refuted this claim (e.g. Samuelsson and Ahlin, 1976; Gorbatschev and ˚ ha¨ll et al., 1997b). Considering the Welin, 1980; A widespread magmatic activity at c. 1.4 Ga over most of southern Sweden, local, but abundant, charnockitization and direct evidence for elevated metamorphic conditions (zircon growth in the mafic gneiss sample and regional formation of granitic neosome in the gneisses; this study), this episode must have been characterized by high heat flow and consequently a weakened crust. Ductile, intra-cratonic deformation, confined to discrete zones or belts, seems possible, if not likely, although no unequivocal evidence has been reported. In particular, the late Hallandian emplacement of large megacrystic granites at

1.38 Ga must have caused deformation in adjacent gneisses as these dry intrusions intruded during a period of weakened crust. 4.3. Sveconorwegian events D3 caused a dominantly linear fabric of variable regional intensity. In spite of this variability, it is typically easily distinguished as the first fabric in the Group 2 dykes ( Table 1). In the study area, it is bracketed only by the 1.40 Ga ages and the post-kinematic 946+6/−4 Ma pegmatite, but is interpreted as Sveconorwegian ( Table 1). Such inferences are supported from other areas where it affects rocks as young as 1224 Ma (Berglund and Connelly, 1994). The post-kinematic pegmatite intruded at 946+6/−4 Ma constrains the last component of penetrative ductile deformation (D4). The pegmatite is coeval with other late¨ tran terrane, orogenic features elsewhere in the A including granitic dykes (Mo¨ller and So¨derlund, 1997), titanite cooling ages (Connelly et al., 1996; Wang et al., 1998) and Ar–Ar hornblende ages (Page et al., 1996a; Wang et al., 1996), that collectively imply that peak Sveconorwegian conditions were replaced by late-stage extension and cooling by this time. Extensive Sveconorwegian recrystallization during upper-amphibolite to granulite facies conditions has been well documented by mineral assemblages and equilibria (e.g. Johansson et al., 1991, 1998; So¨derlund, 1996; Mo¨ller and So¨derlund, 1997; this study). The lack of migmatization within and along the contacts of the 1426 and 1399 Ma Group 2 dykes (Fig. 3) demonstrates that the main migmatitic banding that they cut in the host gneisses in the Varberg–Halmstad area cannot be Sveconorwegian. However, Sveconorwegian migmatization does occur within shear zones and is easily demonstrated in marginal parts of the 1.38 Ma Torpa granite (e.g. Talbot and Heeroma, ˚ ha¨ll et al., 1997b) and the c. 1.37 Ga 1989; A Tja¨rnesjo¨ granite (Andersson, 1997; Mo¨ller and So¨derlund, 1997). Major Sveconorwegian shear zones containing Sveconorwegian migmatites have also been recognized in gneisses well within the ¨ tran terrane (e.g. Mo¨ller and So¨derlund, 1997). A An explanation for the surprising lack of a regional migmatitic overprint, considering the high

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Sveconorwegian temperatures and regional D3 lineation in the study area, may lie in earlier dehydration during the Gothian and Hallandian events. This general lack of regional Sveconorwegian migmatitic overprinting, and the extensive static recystallization of all pre-Sveconorwegian assemblages, are compatible with the tectonic models in which ¨ tran terrane represents a distal segment lying the A east of the main axis of the Sveconorwegian orogen ˚ ha¨ll, 1996; A ˚ ha¨ll et al., 1997a). If correct, this (A segment would have been buried and heated during the (south)eastward thrusting (e.g. Berthelsen, 1980; Park et al., 1991; Stephens et al., 1996), a model consistent with high-temperature overprinting during a clockwise P–T–t loop (Mo¨ller and So¨derlund, 1997; Johansson et al., 1998; Mo¨ller, 1998). Migmitization within Sveconorwegian shear zones may have been encouraged by fluid flow. The late Sveconorwegian cooling history in the Steninge area can be estimated using the ages and estimated closure temperatures for monazite, titanite and rutile (720°C for monazite — Copeland et al., 1988; Parrish, 1990; Mezger et al., 1991; 650°C for titanite — Tucker et al., 1987; Heaman and Parrish, 1991; Mezger et al., 1991, 1993; and 380–420°C for rutile — Mezger et al., 1989). Following initially fast exhumation rates immediately after peak Sveconorwegian metamorphism (Mo¨ller and So¨derlund, 1997; Mo¨ller, 1998), cool¨ tran terrane apparently slowed to ing of the A approximately 5–11°C/m.y. between 948 and 932 Ma, and 2.5–5°C/m.y. between 932 and 878 Ma, depending upon the choice of closure temperatures. The slow cooling rates over this period [analogous to those documented by Mezger et al. (1991) for the Adirondacks] are compatible with unroofing during erosion and isostatic uplift.

5. Summary Field relationships and U–Pb data for eight samples from the Varberg–Halmstad region constrain the polymetamorphic evolution in an area of high-grade Sveconorwegian overprinting in the ¨ tran terrane, SW Sweden. The 1664±7 Ma proA tolith age for the Sa¨rdal orthogneiss coincides with ages for other migmatized granitoid intrusions (Connelly et al., 1996; Johansson, 1998) and thus

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further documents an important magmatic event ¨ tran terrane. This event coincided with the in the A N–S trending TIB 3 magmatism to the east. Metamorphic zircon growth in mafic Steninge dykes at 1654±9 Ma is related to Gothian, D2 amphibolite-facies deformation. The earlier D1 event is constrained between 1664±7 and 1654±9 Ma. Given the small amount of time between D1 and D2, they may be part of a continuous event that included the c. 1.66–1.65 Ga emplacement of the Group 1 mafic Steninge dykes. These data constrain the gneissic D1/D2 banding in this region to a 1664–1654 Ma interval in the Gothian orogeny. Two abundant rock units, a 1426+9/−4 Ma suite of irregular granitic dykes and a straightwalled 1409–1399 Ma suite of pegmatites (of which one has been dated here at 1399+7/−6 Ma), were emplaced after the major gneiss-forming event(s) in the area. This evidence is consistent with that ¨ tran terrane and requires from other parts of the A that the major regional gneiss-forming event(s) in ¨ tran terrane was pre-1.43 Ga. the A Post-1400 Ma deformation is regarded as Sveconorwegian and includes a regional fabricforming event (D3), thorough mineral recrystallization and local folding (D4). A post-kinematic pegmatite is dated at 946+6/−4 Ma and constrains that last phase of Sveconorwegian deformation (D4) to 946 Ma. The new data support previous inferences of a high-temperature Hallandian event by establishing a protracted thermo-magmatic episode between 1.44 and 1.38 Ga. This thermo-magmatic episode is characterized by widespread and varied magmatic activity over most of southern Sweden, but a number of features are unique to the Varberg– Halmstad region: (1) growth of metamorphic zircon, here dated at 1438+12/−8 Ma in a mafic gneiss, (2) regional formation of 1426+9/−4 Ma granitic neosomes and minor dykes, (3) crystallization of the partly charnockitic Varberg CGA (a sample of which yields a 1399+12/−10 Ma age), and (4) local but widespread charnockitization in the gneisses. All data collectively indicate a 1.44– 1.38 Ga Hallandian episode that was characterized by high heat flow. Although no coeval deformation has yet been identified, the late Hallandian emplacement of large, dry granitic intrusions (at

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c. 1.38 Ga) is predicted to have caused significant ductile deformation in adjacent gneisses. Static recrystallization textures in the post-kinematic pegmatite and mafic dykes indicate that thermal conditions remained elevated after 946+6/−4 Ma. Metamorphic monazite (948± 4 Ma), titanite (935±5 Ma and 932±7 Ma) and rutile (879±10 Ma) also indicate that temperatures in the study area remained elevated until c. 879 Ma. The progressive younging of these mineral ages, combined with their closure temperatures, reflects slow cooling rates after 946 Ma, consistent ¨ tran terrane due with isostatic unroofing of the A to post-contractional erosion. The Sm–Nd mineral isochron age of 878±23 Ma from a mafic gneiss coincides with other Sm–Nd results from the area and must represent post-peak metamorphic re-equilibration of Nd.

Acknowledgements Special thanks to Benjamin Pursell, Lena Lunqvist and Sven Larsson for their assistance during the field component of this research. Kathryn Manser is gratefully acknowledged for her technical skills in the U–Pb geochronology laboratory, as is Inger Lundqvist (SGU, Go¨teborg) for support of the geochronology program. Comments on C.A.C.’s MSc. thesis, upon which this paper is based, by Sharon Mosher, Leon Long and Laurie Schuur are appreciated. Journal reviews by Klaus Mezger and Ulf So¨derlund were very helpful. This study was supported by grants from the Geological Society of America (C.A.C.), Sigma Xi (C.A.C.), The University of Texas at Austin Geology Foundation (C.A.C. and J.N.C.) and from the Swedish Natural ˚) Research Council, grant G-GU 10286-304 ( KIA and Geological Survey of Sweden, grant SGU ˚ ). 03-826/93 ( KIA

at Austin. Mineral fractions were carefully selected for U–Pb geochronology on the basis of optical characteristics using binocular and petrographic microscopes. Subsequent characterization of internal structures utilized cathodoluminescence (CL) microscopy. All fractions were air-abraded using the technique of Krogh (1982), then re-examined and cleaned with distilled 4 N HNO , quartz-distilled 3 water and distilled acetone. Grains were weighed into Teflon dissolution bombs with a mixed 205Pb/235U isotopic tracer and dissolved with HF and HNO . For zircon analyses, U and Pb were 3 chemically separated using minicolumns (55 ml resin volume) following the procedure of Krogh (1973). Pb and U blanks were estimated to be 1 and 0.25 pg, respectively. For titanite analysis, HBr chemistry employed 120 ml columns with a procedural blank of 3 and 0.25 pg for U and Pb, respectively. U and Pb were loaded together onto outgassed, zone-refined Re filaments using silica gel and phosphoric acid. Isotopic analyses of larger samples were performed on a Finnigan MAT 261 thermal ionization mass spectrometer in static multicollection mode with 204Pb measured in the axial secondary electron multiplier (SEM ) ion counter. Smaller samples were measured in peak jumping mode using only the SEM – ion counter. Ages were calculated using an unpublished program written by J. Connelly incorporating the decay constants of Jaffey et al. (1971). Analytical errors of isotope ratios, reported at two sigma, were calculated with a program modified by J. Connelly after an unpublished error propagation program written by L. Heaman ( University of Alberta, Edmonton, Canada). Linear regressions were calculated using the procedure of Davis (1982); an acceptable probability of fit is 10%, corresponding to an MSWD of 2.

Appendix B: Sm–Nd techniques Appendix A: U–Pb techniques Eight samples of approximately 10 kg were crushed for geochronology using standard mineral separation techniques at The University of Texas

Sm–Nd analyses for apatite, clinopyroxene and titanite were performed at The University of Texas at Austin. Fractions were hand-picked to eliminate grains with inclusions and fractures. Apatite was

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cleaned with quartz-distilled water and distilled acetone; clinopyroxene and titanite were cleaned with 2 N nitric acid, quartz-distilled water and distilled acetone. Samples were put into Teflon capsules with a mixed 149Sm/150Nd isotopic tracer and dissolved using appropriate acids (HF and HNO for silicates, HCl for phosphates). The 3 REEs were isolated from dissolved apatite and titanite, followed by 8 cm×0.7 cm HDEHP columns to separate Nd and Sm. Chemical separations for clinopyroxene required an additional step employing cation columns to remove Fe prior to REE-Spec chemistry. The resulting procedural blank was 25 pg for both Nd and Sm. Nd and Sm were loaded separately onto outgassed, rhenium filaments using phosphoric acid and 3 N HCl. Analyses were carried out on a Finnigan MAT 261 thermal ionization mass spectrometer in multidynamic mode. 146Nd/144Nd ratios were normalized to 0.7219. Two Nd standards were analysed at UT during the period of these analyses: (1) the Ames Nd Standard yielded an average 148Nd/144Nd value of 0.512088±8 (n=6), and (2) the CIT Nd Standard yielded an average 148Nd/144Nd value of 0.511899±6 (n=11). All errors are reported at two sigma. Isochron ages were calculated using ISOPLOT written by K.R. Ludwig (Ludwig, 1990), incorporating decay constants of Lugmair and Marti (1978).

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