Download - Zircon geochronology in polymetamorphic gneisses in the Sveconorwegian orogen, SW Sweden: ion microprobe evidence for 1.46–1.42 and 0.98–0.96 Ga reworking

Precambrian Research 113 (2002) 193–225

Zircon geochronology in polymetamorphic gneisses in theSveconorwegian orogen, SW Sweden: ion microprobeevidence for 1.46–1.42 and 0.98–0.96 Ga reworking

Ulf Soderlund a,*, Charlotte Moller a, Jenny Andersson a, Leif Johansson a,Martin Whitehouse b

a Department of Mineralogy and Petrology, Institute of Geology, Uni�ersity of Lund, Sol�egatan 13, 22362 Lund, Swedenb Swedish Museum of Natural History, S-104 05 Stockholm, Sweden

Received 7 January 2000; accepted 2 July 2001

Abstract

Ion microprobe U–Th–Pb analyses of zircons in variably metamorphosed and veined orthogneisses in the southernpart of the parautochthonous Eastern Segment of the Sveconorwegian (1.20–0.90 Ga) orogen, SW Sweden, broadlydefine two age groups, oscillatory and sector zoned magmatic zircon cores yield 1.70–1.68 Ga while overgrowths,homogeneous crystals, and recrystallized domains in primary zircon yield 1.46–1.42 Ga. In addition, a late-kinematicpegmatite was dated at 0.96 Ga, while a penetratively deformed granite dyke contains both 1444�8 Ma magmaticand 982�15 Ma metamorphic zircons. The 1.70–1.68 Ga ages date the orthogneiss protoliths and fall in the sameage range as well-preserved rocks of the Transscandinavian Igneous Belt that forms a major part of the crust east ofthe Sveconorwegian orogen. Despite Sveconorwegian penetrative deformation under granulite to upper amphiboliteconditions, secondary zircons yielding Sveconorwegian ages are virtually absent in the 1.70–1.68 Ga orthogneisses butare abundant in rocks younger than ca. 1.45 Ga. It is suggested that Zr hosted in magmatic phases was redistributedto form new zircon during the 1.46–1.42 Ga event, resulting in a mineralogy in which the main minerals weredepleted in Zr. These data, therefore, imply that high-grade metamorphism may occur without associated growth ofnew zircon. Furthermore, the absence of secondary zircons with ages �ca. 1.46 Ga suggests a re-assessment ofmodels calling for extensive Gothian deformation and metamorphism in the Eastern Segment. © 2002 ElsevierScience B.V. All rights reserved.

Keywords: 1.4 Ga; Sveconorwegian; Hallandian; Gothian; Ion probe; Metamorphic zircon; Sweden

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1. Introduction

One approach to gain information about theearly history in geologically complex regions is toinvestigate zircon populations in the oldestlithologies. The high-grade, southern part of the

* Corresponding author. Fax: +46-46-121477.E-mail address: [email protected] (U. Soderlund).

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U. Soderlund et al. / Precambrian Research 113 (2002) 193–225194

Eastern Segment of the Sveconorwegian orogen inSW Scandinavia (Fig. 1) has traditionally beenregarded as polymetamorphic, metamorphosedand deformed during both the accretional Goth-ian (ca. 1.75–1.55 Ga) and the collision-relatedSveconorwegian (ca. 1.20–0.90 Ga) orogenies(e.g. Gorbatschev and Bogdanova, 1993). SinceSveconorwegian peak metamorphic temperaturesexceeded 700 °C, Sm–Nd and 40Ar–39Ar mineral,

and U–Pb titanite, generally yield late Sveconor-wegian ages most of which represent cooling (e.g.Johansson et al., 1991; Wang and Lindh, 1996;Connelly et al., 1996; Wang et al., 1996, 1998;Christoffel et al., 1999). Therefore, these age re-sults cannot be used to discriminate between Sve-conorwegian and older tectonic events.

In this study, we have used back-scatter elec-tron (BSE) imaging and ion microprobe U–Th–

Fig. 1. Tectonic sketch map of SW Sweden modified after Wahlgren et al. (1994). The framed area depicts the location of Fig. 3.

U. Soderlund et al. / Precambrian Research 113 (2002) 193–225 195

Pb analysis to characterize and date igneous crys-tallization and later zircon growth and reworkingin the oldest rock units. Five orthogneiss varieties,a deformed granite dyke and a late-kinematicpegmatite, were sampled from the southern, inter-nal part of the Eastern Segment. Different typesof zircon are distinguished by their internal mor-phology, Th–U chemistry and U–Pb ages, andare related, as far as possible, to magmatic andmetamorphic processes. The results allow us toevaluate the controversial tectonic history of theEastern Segment. From the distribution of sec-ondary zircon in different generations of mag-matic rocks we also discuss the ability of zircon torecord metamorphism.

2. Tectonic setting of the Eastern Segment

2.1. Regional setting and lithologies

The Precambrian of southern Sweden belongsto two major orogens, the Palaeoproterozoic(1.89–1.86 Ga) Svecokarelian orogen in the east,and the Meso- to Neoproterozoic (1.20–0.90 Ga)Sveconorwegian orogen in the west (Fig. 1, Gor-batschev and Bogdanova, 1993). The Svecokare-lian orogen is intruded by the ca. 1.85–1.65 GaTransscandinavian Igneous Belt (TIB), a large,coherent, N–S trending belt, partly affected bySveconorwegian reworking in the west (e.g. Lar-son et al., 1998; Soderlund et al., 1999).

The Sveconorwegian orogen is made up oflarge lithotectonic units separated by N–S trend-ing zones of ductile deformation (Berthelsen,1980; Park et al., 1991; Wahlgren et al., 1994;Berglund, 1997). The parautochthonous EasternSegment is delimited to the east, north of LakeVanern, by the Sveconorwegian Frontal Deforma-tion Zone (SFDZ, Wahlgren et al., 1994). Furtherto the south, the location of the SFDZ is morespeculative and is only tentatively illustrated inFig. 1. In the west, the west-dipping MyloniteZone bounds the Eastern Segment from the Ide-fjorden terrane (Fig. 1). In its northern and cen-tral parts, the Mylonite Zone has accommodatedSveconorwegian transpressional deformation,with left-lateral and top to the east movements

under retrogressive metamorphic conditions(Stephens et al., 1996). South of Lake Vanern, theyoungest deformation was extensional, most likelyassociated with exhumation of the high-gradesouthern part of the Eastern Segment (Berglund,1997). In this area, the Mylonite Zone marks amajor contrast in metamorphic grade, separatingSveconorwegian eclogites and high-pressure gran-ulites in the Eastern Segment from epidote-amphi-bolite facies rocks in the Idefjorden terrane.

North of Lake Vanern, both TIB rocks and ca.1.57 Ga mafic dykes are deformed, semi-penetra-tively in the east and penetratively in the west(Wahlgren et al., 1994; Soderlund et al., 1999).Also south of Lake Vanern, progressively de-formed and metamorphosed TIB rocks can befollowed from east toward the west (Samuelssonet al., 1988). Veined orthogneisses in the Ulrice-hamn–Boras area have yielded U-Pb zircon agesat 1.70 and 1.69 Ga (Fig. 1, Connelly et al., 1996),suggesting that TIB rocks extend further west intothe interior of the Eastern Segment. Similarlydeformed and veined, but younger TIB rocks,may also exist further west as granitic ca. 1.66 Gaorthogneisses have been recognized in the Halm-stad area (Fig. 1, Johansson, 1998; Christoffel etal., 1999). Granitic to quartzmonzonitic intru-sions, dated at ca. 1.54, 1.46–1.37 and 1.22 Ga(Berglund and Connelly, 1994; Connelly et al.,1996; Lindh, 1996; A� hall et al., 1997; Anderssonet al., 1999, Christoffel et al., 1999), are broadlycoeval with mafic dykes close to the SFDZ in theeast (Johansson and Johansson, 1990; Ask, 1996;Lundqvist, 1996). Still younger mafic dykes (ca.0.95 Ga, Johansson and Johansson, 1990; Soder-lund, unpublished. U–Pb baddeleyite data) be-long to the major Blekinge–Dalarna doleriteswarm (BDD). In southeasternmost Sweden(county of Blekinge), ca. 1.45–1.40 Ga graniteshave intruded ca. 1.77 Ga granitoid gneisses (re-worked TIB rocks, Kornfalt and Vaasjoki, 1999).The continuation of TIB rocks and younger intru-sive rocks across the SFDZ indicates that theEastern Segment rocks developed in an integratedpart of the Fennoscandian Shield rather than inan outboard terrane that was accreted later.

West of the Mylonite Zone, the Idefjorden ter-rane contains 1.62–1.55 Ga granioids and vol-


cano-sedimentary sequences including the ca. 1.64Ga Horred Belt, the ca. 1.62 Ga A� mal Belt andthe 1.60–1.59 Ga Stora Le-Marstrand (SLM)units (A� hall et al., 1995; A� hall and Larson, 2000,and references therein). Thus, the bulk of theIdefjorden rocks appears to be younger thanthose of the Eastern Segment. Structural modelsfor this terrane feature both Gothian and Sve-conorwegian metamorphism, and ductile Gothiandeformation has been dated as late as ca.1.55 Ga(Connelly and A� hall, 1996).

2.2. Tectonometamorphic history of the EasternSegment

North of Lake Vanern, the regional structuresin the Eastern Segment are Sveconorwegian in ageand form a fan-like geometry with gently east-dip-ping foliations in the west that steepen towardsthe east (Wahlgren et al., 1994). The metamorphicgrade increases from east to west reaching amphi-bolite facies. Igneous titanites from penetrativelydeformed ca. 1.7 Ga TIB rocks in the interior partof the segment yield U–Pb ages similar to thezircon ages (1.67–1.66 Ga) whereas a secondarygeneration of titanite dates Sveconorwegian re-working at 976–956 Ma (Soderlund et al., 1999).These data, in agreement with field observations(Larson et al., 1998), suggest that substantialparts of the Eastern Segment north of LakeVanern escaped pre-Sveconorwegian metamor-phism exceeding the closure temperature of theU–Pb titanite system. 40Ar–39Ar homblende agesfall in the interval 1010−960 Ma (Page et al.,1996a) further supporting a thorough Sveconor-wegian overprinting.

South of Lake Vanern, the Eastern Segmentwas reworked under high-grade metamorphicconditions, and at least some parts experienced apolyphase, tectonothermal evolution. In the Ul-ricehamn–Boras area, Sveconorwegian metamor-phism reached upper amphibolite faciesconditions, P–T estimates yield approximately750 °C and 9 kbar (Cornell et al., 1996). U–Pbtitanite analyses are concordant at ca. 950 Ma ordefine Sveconorwegian lower intercept ages (Con-nelly et al., 1996). Conventional U–Pb dating onzircons in leucosome samples also yield Sveconor-

wegian lower intercept ages (Connelly et al.,1996). Based on ion microprobe U–Pb analysis,Cornell et al. (1996) Cornell et al. (1997) datedmetamorphic zircon growth and recrystallisationof primary zircons in the first generation of leuco-some at ca. 970 Ma. A pre-Sveconorwegian tec-tonothermal history in the Ulricehamn–Borasarea is indicated by an upper intercept age of1470�25 Ma, constrained by strongly discordanttitanite analyses (Connelly et al., 1996). At Vrana(see Fig. 1), migmatitic structures are cut by anaplite dyke, dated at 1457�7 Ma (Connelly et al.,1996).

In the southernmost part of the Eastern Seg-ment, Sveconorwegian metamorphism reachedupper amphibolite and high-pressure granulite-fa-cies. P–T estimates for the thermal peak fall inthe range of 680–800 °C and 8–12 kbar (Jo-hansson et al., 1991; Wang and Lindh, 1996), andremnants of Sveconorwegian eclogites have beendocumented (Moller 1998; Moller, 1999). Fig. 2shows the regional distribution of metamorphicand anatectic zircon dated at ca. 980–960 Ma(references given in the figure), similar to the ageof leucosome-hosted zircons in the Ulricehamn–Boras area. U–Pb ages of titanite range betweenca. 945 and 920 Ma (e.g. Wang et al., 1998;Christoffel et al., 1999) and Ar–Ar homblendeages are 1030–920 Ma (Page et al., 1996b; Wanget al., 1996). The older group of hornblende ages(ca. 1030–965 Ma), originally suggested as datingan early Sveconorwegian, crustal thickening event(Page et al., 1996b; Wang et al., 1996), has laterbeen (re-)interpreted as geologically meaningless,carrying a component of excess argon (Wang etal., 1998). Also the Sm–Nd results (ca. 940–880Ma, Johansson et al., 1991; Wang et al., 1996;Christoffel et al., 1999) on upper amphibolite- togranulite-facies assemblages have been re-evalu-ated as these ages cannot date peak metamorphicconditions as first suggested.

In addition to the Vrana locality (Fig. 1), pre-Sveconorwegian migmatization has been demon-strated at a few coastal localities between Varbergand Halmstad. At Steninge (ca. 20 km N ofHalmstad, Fig. 1) a granitic dyke was dated at1426+9/−4 Ma and a pegmatite at 1399+7/−6, both interpreted as postdating regional migma-


Fig. 2. Map of SW Sweden summarizing localities where metamorphic zircon, or zircon related to leucosome formation, haveyielded late Sveconorwegian ages.

tization (Christoffel et al., 1999). Thick ca. 1.43Ga zircon rims in similar orthogneisses nearHalmstad are interpreted as dating this pre-Sve-conorwegian migmatization (Johansson, in prep.)

3. Geologic setting of the Varberg–Ullared area

The Varberg–Ullared area (Fig. 3) is domi-nated by grey plagioclase-rich orthogneisses whichform the country rocks for Sveconorwegian de-compressed eclogites (Moller 1998; Moller, 1999).Reddish granitic gneisses are also common andcover a large area N and NE of Mardaklev (Fig.3). Large, variably deformed bodies of ca. 1.37Ga K-feldspar megacrystic granites of the Tjar-nesjo–Torpa type occur as lens-shaped and dis-rupted km-sized pods (A� hall et al., 1997;Andersson et al., 1999). With few exceptions, allthese more or less gneissic intrusive rocks havebeen completely recrystallized into a high-pressuregranulite or (more commonly) upper amphibolitefacies metamorphic mineralogy, and large por-tions are migmatized or veined. The charnockiticVarberg intrusive, interpreted as co-magmatic

with the Torpa granite, was recently dated at1399+12/−10 Ma (Christoffel et al., 1999).

The eclogite remnants are preserved in less de-formed and less amphibolitized parts of layeredmafic complexes, which occur as lens-shaped bod-ies or semi-continuous sheets up to 1 km thick.Both the precursor of the eclogites and the plagio-clase-rich host gneisses were intruded by 1.44–1.40 Ga granitic and aplitic dykes (Johansson etal., 2001), and thus experienced a common Sve-conorwegian tectonic history. The eclogitizationhas been dated at 972�14 Ma, the age of zirconinclusions in garnet (Johansson et al., 2001). Sub-sequent exhumation and deformation took placeat high-pressure granulite and upper amphibolitefacies conditions (Moller 1998; Moller, 1999), P–T estimates for the thermal peak are 750�45 °Cand 9.5–12 kbar. The rocks are strongly de-formed with a planar gneiss fabric generally asso-ciated with ESE plunging stretching lineationsand the aeromagnetic pattern suggests that thisdeformation has affected at least a 12 km widezone, the Ullared Deformation Zone (UDZ,Moller et al., 1997). The gneiss fabric dips moder-ately and mainly to the NE; sparse kinematic


indicators suggest a down-to-the-east or dextralmovement depending on the orientation of thefoliation (Moller et al., 1997; Moller, 1999).

4. Samples and zircon geochronology

Eight samples for zircon geochronology werecollected from four areas, Gallared S; Gallared N;Dagsas; and Mardaklev (Fig. 3). From the use ofdifferent U–Pb techniques and BSE imaging wepresent protolith ages for gneisses and one de-formed granitic dyke as well as ages for differenttypes of secondary zircon.

Detailed analytical descriptions and criteria forselection of U–Pb ion microprobe data are givenin the Appendix A. U–Pb data of which a major-ity of analyses are discordant were regressed usingthe least squares program of Ludwig (1991a) Lud-wig (1991b), with dates constrained from interceptages in the concordia diagram. Samples domi-nated by concordant analyses are presented asweighted means of individual 207Pb/206Pb ages.

Analyses with 206Pb/204Pb ratios below 8000 and/or those more than 5% discordant are not usedfor age calculation as these properties generallyindicate isotopic disturbance and yield potentiallyerroneous ages. Pb–Pb evaporation ages and cor-responding errors are calculated as the arithmeticmeans of individual data blocks. All ages arequoted at the 2� confidence level.

4.1. Gallared

The rocks in the Gallared area occur in anE–W striking part of the UDZ (Fig. 3). In theplagioclase-rich orthogneisses, primary structures,dykes and veins have been rotated, folded, flat-tened and stretched into a strong E–W strikingfoliation associated with a pronounced subhori-zontal stretching lineation. The orthogneiss vari-eties include, (1) unveined gneiss (common); (2)gneiss with a faint (late-magmatic?) network ofmm-wide white veinlets; (3) gneiss with abundantcm-wide reddish veins without melanosome (com-mon); and (4) gneiss with reddish veins borderedby biotite-enriched melano some.

Fig. 3. Geological map of the Varberg–Ullared area, modified from Moller et al. (1997), and references therein) and Moller (1998).Areas sampled for geochronological studies are marked, DS, Dagsas; GD-N, Gallared North; GD-S, Gallared South; MA� ,Mardakiev.


Fig. 4. Photographs illustrating the field relations and fabrics of the deformed granite-aplite dyke and its plagioclase-rich host gneissat Gallared S. (A) low-angle discordance showing that, prior to deformation, the felsic dyke has crosscut mafic layers (former dykes,xenoliths?), a faint network of mm-scale white veinlets in the plagiclase-rich granitoid and possibly an older gneiss fabric. Arrowsmark the contact between the dyke and the orthogneiss (coin, 2.5 cm, for scale), (B) deformation fabric in the granitic dyke (coin,2 cm, for scale), (C) ribbon structure in the granitic dyke (photomicrograph, crossed polars, width of the photo is �8.5 mm), (D)Penetrative deformation fabric in the plagioclase-rich host gneiss, with anastamosing fabric of alternating clinopyroxene-amphibole-rich (dark) and plagioclase-rich (light) domains (coin, 2 cm, for scale).

4.1.1. Gallared southThe plagioclase-rich orthogneisses at Gallared S

(GD-S in Fig. 3) are dominantly unveined andcontain foliation-parallel, up to 0.5 m thick, maficlayers or dykes as well as reddish to pink apliteand granite dykes. Three samples were investi-gated: a deformed granite dyke, orthogneiss withfaint white veinlets, and dark weakly deformedorthogneiss.

4.1.1.1. Deformed granite dyke (Swedish NationalGrid coordinates, SNG: 633525/132225). At thislocality, a deformed composite granite-aplite dykecross-cuts mafic layers (former dykes?) and a faintnetwork of mm-scale white veinlets in the plagio-clase-rich gneiss (Fig. 4A). The dyke is 2–5 dmthick, E–W striking and dips 70° to the south. Onhorizontal surfaces the dyke is concordant withthe foliation, but on surfaces at high angles to thegently E-plunging stretching lineation a low angle

discordant (ca. 15°) relationship is preserved. Thegranitic parts have a strong deformation fabric,albeit annealed and recrystallized (Fig. 4B), withelongated feldspar domains and ribbons of quartz(Fig. 4C). Local, late stage alterations includesericitization of feldspar, chloritization of garnet,muscovite growth and alteration of biotite byopaques+greenish sheet silicate. The aplitic partsof the dyke have a weak deformation fabric only,which reflects the difficulty to develop and pre-serve pronounced deformation fabrics in this rocktype (fine-grained, poor in Fe–Mg minerals andstrongly annealed). The strong gneissic L–S fabricin the host gneiss is characterized by fine-scaleanastamosing aggregates of alternating clinopy-roxene/amphibole-rich and plagioclase-rich mi-crodomains (Fig. 4D).

Two types of zircon are distinguished optically(Fig. 5A). The first type is colourless and varies inshape from elongated to subhedral. Cracks are


abundant and crystal faces are rare. Their mor-phology indicates different degrees of resorptionof originally prismatic zircons. BSE images showthat only a few crystals of this type have a thin,dark, peripheral margin that possibly representsa metamorphic overgrowth; thus, the variety inshape is probably due to resorption rather thanovergrowth of new zircon. The second type ofzircons are multi-facetted with round to ovalshapes, have high lustre and lack fractures (Fig.5A). They are homogeneous and BSE-darkerthan the first type. The morphological character-istics suggest that the first type of zircon is ig-neous, formed during crystallization of thegranite-pegmatite dyke, whereas the second typeis metamorphic.

Four fractions of subhedral zircon of the firsttype, comprising 2–8 grains in each fraction,were analyzed by the conventional ID-TIMS U-

Pb technique (Table 1). Intercept ages are1443�26 and 88�343 Ma (MSWD=2.6, Fig.5B); the least discordant fractions contain onlytwo grains each. The upper intercept age is in-terpreted as dating the dyke intrusion. Thelower intercept age intersects within error theorigin, indicating variable amounts of modernPb-loss.

To test this interpretation, two similar subhe-dral crystals were selected for Pb–Pb evapora-tion analysis (Kober, 1986, Fig. 5C). Theresulting multi-step 207Pb/206Pb plateau ages of1444�9 and 1443�7 Ma are overlapping(Table 2) such that the accumulated age of1444�8 Ma (2�) constrains the obtained U–Pbage of 1443�26 Ma. The more precise evapora-tion age of 1444�8 Ma is preferred as the ageof the dyke and hence as a minimum age forthe truncated gneissic structures (Fig. 4a).

Fig. 5. Zircon in the deformed granite-aplite dyke, Gallared S. (A) petrographic microscope image showing two different types ofzircon. Subhedral igneous zircons with abundant fractures (left), group of metamorphic, round to oval-shaped, zircons with highlustre (right), (B) U–Pb concordia diagram of zircons in the dyke analyzed by conventional U–Pb technique, (C) Pb–Pbevaporation diagram with age (Ma) versus number of plating steps of igneous zircons (upper plateaus, zircon ZF and ZG) andmetamorphic zircons (lower plateaus, zircon ZH, ZI and ZJ).

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Table 1Conventional U–Pb zircon data for the unveined oithogneiss samples at Gallared S and Dagsas and the deformed granite dyke at Gallared S

Measured ratios 207Pb/206PbConcentrations (ppm)Analysis fraction 206Pb–207Pb–208PbWeight

Pbcom206Pb/204Pba Radiogen. (atom%)b 206Pb/238Ub 207Pb/235Ub Age (Ma)U(�g) Pbrad

Weakly deformed orthogneiss at Gallared79.4–8.0–12.6 0.2688�6 3.712�12175.3 1627�451.0151Z1 (78), 100–150 �m, o.s, cl, 0.2 11280

0.1 2030 77.8–7.9–14.2 0.2827�8 3.972�16 1659�5Z2 (36), 100–150 �m, o.s, cl, abr 23.1 49.3 15.479.8–8.0–12.2 0.684�14 3.721�22 1634�43057198 0.357.3Z3 (10), �100 �m, o.s, cl, i, abr 14.5

235126.6 78.0–8.0–14.0 0.2827�13 3.996�30 1670�11114.3 35.7 0.4Z4 (9), 100–150 �m, o.s, cl, i, abr

Un�eined orthogneiss at Dagsas81.8–8.1–10.1 0.2674�10 3.638�15 1599�33525Z1 (60), �75 �m, o.s, cl 0.757.7183.990

�0.1 17067 82.1–8.0–9.9 0.2651�7 3.579�10 1585�3117Z2 (41), �150 �m, o.s, cl 240.6 67.079.0–7.9–13.1 0.2772�10 3.826�17 1626�6687155.6Z3 (38), �150 �m, o.s, cl, abr 0.252.3173.2

302915.1 75.9–7.8–16.3 0.881�15 4.074�23 1671�4103.3 33.7 0.3Z4 (10), �150 �m, o.s, cl, abrZ5 (20), �100 �m, o.s, cl, abr 78.4–7.9–13.769.5 0.834�11 3.955�17 1645�3144.6 45.0 �0.1 12577

Deformed granite dyke at Gallared S.77.4–7.0-15.5 0.2408�20 3.011�26473.4 1440�5126.77.2Z1 (2), �200 �m, cr, cl 0.3 3693

0.2 8199 78.8–7.2–14.1 0.2419�10 3.035�20 1447�9Z2 (2), �200 �m, cr, cl 9.7 829.1 219.176.7–6.9–16.4 .2205�10 2.743�17 1430�827100.3Z3 (8), �100 �m, cl, abr 19.0 118.7 29.478.4–7.1–14.5Z4 (8), �150 �m, cl, cl 0.2213�1043.7 2.764�14 1438�4191.0 46.4 0.0 13128

Z, zircon; (X), number of crystals or fragments; p, prismatic; o.s, oval shaped; i, inclusions; cr, cracks; cl, colourless; abr, abraded. Crystal sizes are before abrasion.a Corrected for mass fractionation (0.1%�0.04% per a.m.u.).b Corrected for common lead, blank and mass fractionation.


Table 2Pb–Pb evaporation data of zircons from the deformed granite dyke at Gallared S

Number of EvaporationEvaporation Measured isotope ratios Age (Ma)Zirconscansstep temperature(°C)

207Pb/206Pb 206Pb/204Pb Th/U

9Z1 14401 0.09002�46 �100 000 0.17 1423�1019* 1450 0.09118�402 �100 000 0.19 1447�8

3 95* 1460 0.09117�36 �100 000 0.19 1447�84 29* 1470 0.09100�42 �100 000 0.19 1444�9

38* 1470 0.09094�345 �100 000 0.19 1443�710* 14806 0.09079�68 �100 000 0.19 1439�14

200*2–6Age of step 1444�91Z2 20 1450 0.09041�30 �100 000 0.19 1430�62 39* 1460 0.09103�34 �75 000 0.18 1444�7

36* 1460 0.09092�323 �100 000 0.18 1443�729* 1470 0.09093�26 �100 0004 0.18 1443�519* 1470 0.09100�485 �100 000 0.18 1444�10

6 30* 1480 0.09083�34 �100 000 0.18 1440�77 19* 1480 0.09103�44 �100 000 0.18 1445�9

192*2–7Age of step 1443�7392*Z1+Z2 (11 steps) 1444�820 1450 0.07177�461 �30 000Z3 0.059 967�13

2 18* 1460 0.07207�60 �60 000 0.071 981�172Age of step 18* 981�17

28 1450 0.07160�381 �27 000Z4 0.045 960�1120 1460 0.07169�70 �50 0002 0.033 967�2047* 1480 0.07247�423 �26 000 0.031 982�12

3Age of step 47* 982�121Z5 9 1440 0.07287�36 �8500 0.111 964�11

29 1460 0.07386�322 �9000 0.076 990�918 1450 0.07169�52 �27 000 0.053Z6 962�15119* 1450 0.07199�342 �70 000 0.043 980�10

3 19* 1460 0.07248�80 �60 000 0.037 991�234 28* 1460 0.07195�58 �70 000 0.037 977�16

29* 1470 0.07224�485 �40 000 0.043 983�1395*Age of step 982�152–5

160*(6 steps)Z3+Z4+Z6 982�15

*Analysis used in age calculation.

Four zircons with a metamorphic morphology(Fig. 5A) were also analysed using the Pb–Pbevaporation method (Table 2). Grain Z6 yieldedconsistent 207Pb/206Pb ages between the four lastevaporation steps resulting in an age of 982�15Ma (Fig. 5C). This also becomes the accumulatedage when including the last step of Z3 and Z4which provided only a few evaporation-depositsteps (Table 2). The low yield of data from theseyounger zircons is likely due to low concentrationof radiogenic Pb (and U), in accordance with theirdark BSE-colour. The age of 982�15 Ma is

interpreted as dating metamorphism and defor-mation of the granite dyke. The Th/U ratios,calculated from measured 208Pb/206Pb ratios, sup-port the interpretation of two types of zirconshaving typically magmatic (Th/U�0.19) andmetamorphic (Th/U�0.04) origins (Table 2, cf.Williams and Claesson, 1987).

4.1.1.2. Orthogneiss with white �einlets. The greyorthogneiss at Gallared S, sampled �2 dm fromthe dated granite dyke, lacks coarse veins but hasfaint, mm-scale, white veinlets (Fig. 4A and D).


The L–S fabric is defined by anastomosing darkand light micro-domains made up of fine-grainedgranoblastic and anhedral minerals. Dark do-mains are dominated by clinopyroxene, plagio-clase, homblende, garnet, quartz, scapolite andopaques. Light domains are made up by aggre-gates of feldspars, mainly antiperthite, and larger,elongated grains of quartz. Local late stage alter-ation includes sericitization of feldspar and scapo-lite and fine-grained alterations of pyroxene.Mafic dykes at the same outcrop are foliatedamphibolites with fine-grained anhedral garnet,pyroxene and scapolite.

The majority of zircons are complex with inneroscillatory zoned volumes that often show signs of

resorption, e.g. have smooth, sometimes undulat-ing, outlines. Some zircons show sector zoningtypical for granitoid zircons (Vavra, 1994), withstraight or smoothly curved boundaries betweenadjacent sectors (Fig. 6A). The crystals have ho-mogeneous rims, most of which are distinct over-growths, which appear light in BSE-images (Fig.6A and B). The thickness of the overgrowthsvaries around a single crystal, which is the maincause of their subhedral shapes and varyingmorphologies.

The oscillatory and sector zoned zircons areinterpreted to be igneous, grown during crystal-lization of the protolith. Ion probe results fromfive discordant grains define an upper intercept at

Fig. 6. (A) and (B) are back-scatter electron images of characteristic zircons from the orthogneiss at Gallared S. Abbreviations usedhere and elsewhere: oz, oscillatory zoning; sz, sector zoning; pgb, planar growth band; dz, diffuse zoning; bs, bright zircon; gs, ghoststructures; og, overgrowth; rp, signs of resorption; ftz, fir-tree zoning. Round circles, center of ion-probe spot with corresponding207Pb/206Pb age. Approximate spot size is 25 �m. (A) BSE-image of zircon with a relatively BSE-dark, sector-zoned, inner part anda lighter, oscillatory-zoned, outer part; both belonging to the same generation around 1.7 Ga. The rounded outline of thesector-zoned inner part suggest that zircon growth from the magma was not continuous, interrupted by events of resorption. A thinhomogeneous BSE-light rim, interpreted as an overgrowth, can be seen along the right side of the zircon, (B) oscillatory zoned zirconwith a light overgrowth to the left. Marked ellipses correspond to n215-07a and n215-12a in Table 3, (C) ion mircoprobe data ofzircons from the orthogneiss at Gallared S, (D) ion microprobe data of BSE-light zircon overgrowths.


Fig. 7. [U] versus Th/U diagram for zircons in unveinedorthogneiss samples at Gallared S, and vein sample and late-kinematic pegmatite at Gallared N. Overgrowths in the or-thogneiss sample at Gallared S are marked G.

Fig. 8. U–Pb concordia diagram showing conventional U–Pbanalyses of zircons from the unveined orthogneiss at GallaredSouth.

hedral minerals; dark micro-domains are domi-nated by clinopyroxene, plagioclase, garnet,homblende, biotite, orthopyroxene, scapolite andopaques, and light microdomains by antiperthiticfeldspars and quartz. Late stage alteration is veryminor but include fine-grained dusty alterationsalong cleavage planes and rims of pyroxene andscapolite.

This orthogneiss variety, sampled for conven-tional U–Pb dating, contains colourless andtransparent zircons with elongated to prismaticcrystal shapes. Although three of four zircon frac-tions were extensively abraded, analyses plot dis-cordantly with intercept ages of 1698�109 and825�485 Ma (MSWD=4.5, Fig. 8). The discor-dance and poor precision in ages are interpretedto reflect zircon complexity with ca. 1.7 and 1.45Ga old zircon components weakly overprinted bymodern Pb-loss, as is demonstrated by the ionmicroprobe results for the first gneiss sample fromGallared S (see above).

4.1.2. Gallared North (SNG: 633880/131975)At Gallared N (GD-N in Fig. 3), field relations

are similar to Gallared S, but the gneisses are hereextensively veined and migmatized (Fig. 9A andB). Late-kinematic, pinch-and-swell shaped peg-matitic segregations are foliation-parallel, rangingfrom 2 to 15 cm in thickness (Fig. 9C). 600 msouth of the sampled outcrop, amphibolitized anddeformed remnants of kyanite eclogites occur

1698�12 Ma (MSWD=0.6, Fig. 6C) whereastwo other grains gave markedly younger 207Pb/206Pb ages (Table 3). The zircons are consistentlylow in [U] and have Th/U-ratios around 0.8 (Fig.7). Three analyses of light homogenous rims yield1464�8 Ma (MSWD=0.4, Fig. 6D). Th/U ra-tios are highly variably ranging between 0.04 and0.93 (Table 3). A fourth analysis (n215-03a inTable 3) is significantly younger (1422�14 Ma)and lowers the age to 1458�33 (MSWD=12) ifincluded. This analysis is rejected despite plottingconcordantly within error, having high 206Pb/204Pb ratio and the preferred age result is basedon only three analyses. However, concidering theage results of zircons in other samples (see Fig. 17and Appendix A, ‘procedures of data selection’)we think it is realistic to suspect some Pb-loss forthis analysis.

4.1.1.3. Weakly deformed orthogneiss (SNG:633475/132175). Approximately 600 m SW of thedated granite dyke, the plagioclase-rich granitoidis only weakly deformed. In the centre of darkmicro-domains, strained megacrysts of clinopy-roxene with fine exsolution-lamellae of orthopy-roxene are preserved; these crystals are probablyremnants of primary magmatic clinopyroxene.Most of the rock is, however, recrystallized intofine-grained aggregates of granoblastic or an-

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Table 3Ion probe U–Pb data, Ga1lared S and N. D=Discordance (positive values indicates normal, negative values indicates reverse)

PbU 206Pb/204PbGrain spot 206Pb/238UTh Th/U 207Pb/235U 207Pb/206Pb D206Pb/238U 207Pb/206Pb207Pb/235Uage�2�(Ma) (%)�1� (%) (Ma)(ppm)(ppm) �1� (%)�1� (%)(ppm)

Gallared S, orthogneiss with white �einlets

Oscillator� or sector-zoned zircons0.3221 (1.3) 4.630 (1.4) 0.1043 (0.38) 1800276 17551.1 1702�14 −4126280n215-01a* �105

21 7670 0.2769 (2.9) 3.718 (3.0) 0.0974 (0.93) 1576 1575 1575�34 −435n215-04a 0.58600.2549 (2.3) 3.382 (2.5) 0.0962 (0.95)n215-04b 146474 1500 1552�36 +245 0.58 24 �105

0.2569 (2.2) 3.534 (2.2) 0.1000 (0.54) 1474 1535 1620�2041300 +60.81 38n215-06a 109 9184300165 0.2768 (1.5) 3.883 (1.5) 0.1017 (0.43) 1575 1610 1656�16 +385 0.53 58n215-06b

0.2797 (3.3) 4.002 (3.5) 0.1034 (1.3) 1590 1635 1693�46n215-07a* +169 50 0.66 25 603000.2885 (2.1) 4.122 (2.1) 0.1036 (0.51) 1634 1659 1690�18�105 0128n215-08a* 510.93126

58600107 0.3182 (2.1) 4.619 (2.3) 0.1053 (91) 1781 1753 1719�34 053 0.58 43n215-10a*0.3176 (1.5)n215-11a* 4.597 (1.6)186 0.1050 (0.57) 1778 1749 1714�20 −1130 0.76 78 �105

O�ergrowths0.2475 (3.6) 3.118 (3.7) 0.0914 (0.59) 1425116 14370.93 1455�22 040119n215-02a* 67100

92 30300 0.2495 (1.1) 3.090 (1.2) 0.0898 (0.36) 1436 1430 1422�14 041n215-03a 0.123290.2382 (1.5) 3.007 (1.6) 0.0916 (0.35) 1377 1410 1459�14 +3n215-05a* 305 11 0.04 80 440000.2694 (0.68) 3.422 (.72) 0.0921 (0.23) 1536 1510 1470�859100 −4n215-l2a* 147490 35 0.08

Gallared N, �ein sample

Oscillator�-zoned cores57 �105 0.3194 (1.8) 4.594 (1.9) 0.1043 (0.46) 1787 1748 1703�16 −2n217-01a* 88137 0.6671 10100 0.2607 (1.6) 3.621(1.7) 0.1007 (0.46) 1493 1554 1638�18 +7n217-05a 210 139 0.58

0.3012 (3.1) 4.148 (2.2) 0.1000 (0.55) 1697 1664 1622�205680 −10.52 38n217-08a 100 5112800307 0.2860 (1.4) 4.039 (1.5) 0.1024 (0.39) 1621 1642 1669�14 0168 0.52 111n217-09b*66900164 0.2881 (1.6) 4.063 (1.7) 0.1023 (0.62) 1632 1647 1666�22 0141 0.87 64n217-13b*

0.2950 (1.2) 4.172 (1.3) 0.1026 (0.50) 1667 1669 1671�18�105 0n217-18a* 870.93206215111 22200 0.2385 (1.6) 3.044 (1.6) 0.0926 (0.35) 1379 1419 1479�14 +5n217-12b 30417 0.05

Back-scatter bright zircon0.2566 (1.5) 3.256 (1.6) 0.0920 (0.27) 14729 1471n217-02a* 1468�10 0906 0.008 255 74000

425 �105 0.2472 (1.0) 3.124 (1.0) 0.0917 (0.17) 1424 1439 1461�6 +19n217-06a* 0.00515720.2537 (.84) 3.147 (.86) 0.0900 (0.20) 1457n217-04b* 14441556 1425�8 −115 0.010 432 597000.2627 (1.4) 3.296 (1.4) 0.0910 (0.15) 1504 1480 1446�6�105 −21140n217-09a* 3280.0089

�105797 0.2538 (1.8) 3.192(1.8) 0.0912 (0.28) 1458 1455 1451�10 06 0.007 221n217-10a*�1051256 0.2538 (1.0) 3.154 (1.1) 0.0901 (0.28) 1458 1446 1428�10 07 0.005 348n217-11a*

0.2373 (1.6) 2.956 (1.7) 0.0904 (0.58) 1373 1396 1433�228180 +2n217-12a* 2890.03391102n217-15a* 0.2542 (2.4)1477 3.203 (2.4) 0.0914 (0.35) 1460 1454 1454�14 06 0.003 410 �105

Homogeneous o�ergrowths170 61200 0.2842 (1.2) 3.557 (1.2) 0.0908 (0.22) 1613 1540 1442�8 −1185n217-04a 0.19523

0.2181 (1.7) 2.704 (1.8) 0.0900 (0.37) 1272 1329 1423�1412100 +9n217-04c 2580.089410550.2574 (2.5) 3.232 (2.6) 0.0911 (0.74) 1477n217-06b* 146579 1448�28 079 1.1 28 685000.2724 (1.1) 3.431 (1.3) 0.0914 (0.64) 1553 1512 1454�24�105 −5170n217-07a* 610.84126

17700279 0.2482 (1.5) 3.019 (1.6) 0.0882 (0.38) 1429 1412 1387�14 080 0.26 81n217-11b0.2233 (1.7) 2.589 (2.1) 0.0841 (1.1) 1299 1298 1295�44n217-13a 0126 79 0.59 36 22500.2378 (2.8) 2.919 (2.8) 0.0891 (0.64) 1373 1387 1405�2433700 0132n217-14a 420.81115

39500148 0.2562 (2.3) 3.242 (2.4) 0.0918 (0.72) 1471 1467 1462�28 0110 0.72 49n217-14b*n217-14c 0.2368 (1.7)539 2.881 (1.9) 0.0882 (0.69) 1370 1377 1387�26 0120 0.21 147 65900

0.2379 (1.4) 2.886 (2.1) 0.0880 (1.6) 1376 1378 1382�60 08800.3178266n217-17a 78

Homogenous (non-complex) zircon36 �105 0.2649 (3.6) 3.303 (3.7) 0.0904 (0.72) 1515 1482 1435�28 0101n217-03a* 1.19740 56400 0.2510 (2.8) 3.170 (2.9) 0.0916 (0.74) 1443 1450 1459�28 0129n217-19a* 1.2110

0.2387 (1.7) 3.086 (2.2) 0.0938 (1.4) 1380 1429 1504�521520 +6n217-20a 491.4212131

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Table 3 (Continued)

U Th/U Pb 206Pb/238U 207Pb/235U 207Pb/206Pb 206Pb/238U 207Pb/235U 207Pb/206Pb DThGrain spot 206Pb/204Pb(Ma)(ppm) age�2��1� (%) (%)(ppm) �1� (%)�1� (%) (Ma)(ppm)

51 �105 0.2461 (1.4) 3.090 (1.6) 0.0911 (0.68) 1418 1430 1448�26 0n217-21a* 152 132 0.880.2480 (2.2) 3.123 (2.2) 0.0913 (0.54) 1428 1438 1453�20 0�105n229-01a* 137 84 0.56 43

Gallared N, late-kinematic pegmatite

Core0.2779 (1.0) 3.534 (1.1) 0.0922 (25) 1581 15350.045 1472�10417 −622 128 78600n218-01b

462 670 0.2137 (1.2) 2.667 (1.4) 0.0905 (0.77) 1249 1319 1436�30 +12163 0.30n218-07a 125�1051653 0.2485 (1.0) 3.107 (1.1) 0.0907 (14) 1431 1434 1440�6 0407 0.23 476n218-09a

Homogeneous o�ergrowths0.1793 (1.6) 1.759 (1.7) 0.0711 (0.72) 10634 10300.036 961�30 −8118n218-01a 23 14030

129 88300 0.1675 (1.1) 1.646 (1.1) 0.0713 (0.27) 998 988 966�12 −118n218-02a* 0.0277110.1690 (1.7) 1.651 (1.8) 0.0709 (0.36) 1007 990 954�14�105 −2520n218-03a* 960.05726

�105592 0.1692 (0.81) 1.654 (0.84) 0.0709 (0.25) 1008 991 955�10 −417 0.030 108n218-04a*25200237 0.1497 (2.8) 1.442 (3.1) 0.0699 (1.4) 899 907 925�56 08 0.030 38n218-05a*

0.1538 (0.88) 1.497 (.93) 0.0706 (0.32) 922 929 946�1481800 +1n218-06a* 1460.02827878n218-08a 0.1597 (1.7)161 1.589 (1.8) 0.0722 (0.61) 955 966 991�24 +15 0.026 28 83100

0.1549 (1.9) 1.517 (2.0) 0.0710 (0.66) 929 937 958�26�105 029n218-09b* 170 6 0.0320.1646 (0.92) 1.547 (1.0) 0.0682 (0.39) 982 949 874�16 −11n218-01a 858 26 0.031 152 280000.1590 (1.5) 1.571 (1.5) 0.0717 (0.42) 951 959 977�18 095200n230-01a* 442 28 0.055 77



Fig. 9. Photographs illustrating the field relations and fabrics of veined gneiss and associated rocks at Gallared N. (A)Plagioclase-rich gneiss extensively veined by reddish granitic material. Note the strong overprinting deformation (coin, 2.5 cm, forscale), (B) reddish vein bordered by biotite-rich melanosome. Note the overprinting deformation fabric and the small refractionangle of the foliation in the vein and the host rock (width of the photograph is �15 cm), (C) late-kinematic, pinch and swell shapedpegmatitic segregation. Arrows and white line outline the contact between the segretate and orthogneiss (coin, 2.5 cm, for scale). (D)former kyanite eclogite, now extensively amphibolitized and veined. The lens-shaped domain in the centre preserve pseudomorphsof plagioclase+sapphirine (white dots) after kyanite (coin, 2.5 cm, for scale).

(Fig. 9D). Two samples were investigated, agranitic vein (Fig. 9B) and a late-kinematic peg-matite (Fig. 9C).

4.1.2.1. Vein sample. The veins at Gallared N aregenerally a few cm thick and reddish in colour,locally bordered by a thin biotite-rich melanosome(Fig. 9B). They are tightly to isoclinally folded andstretched, and accentuate the gneissic fabric in therock. The fabric is associated with a mineralorientation and a strong ESE-trending and gentlyplunging stretching lineation. The sampled vein haslens-shaped, elongated quartz-domains set in afeldspar-rich matrix with plagioclase (partly anti-perthitic) and microcline with minor myrmekite.The host rock has ribbon quartz and stronglyelongated dark aggregates in a feldspar-rich matrix.Mafic minerals are homblende, garnet, biotite andopaques.

Most zircons are complex with oscillatory and/orsector zoned cores surrounded by BSE-dark over-growths that are homogeneous or have faint planargrowth bands (Fig. 10). Some of the magmaticallyzoned cores are rounded, suggesting resorptionprior to overgrowth. A striking feature is the occur-rence of BSE-bright domains which often truncateboth the primary growth bands in the core and thecore-overgrowth boundary (Fig. 10A). In somecrystals the entire core is replaced by BSE-bright,relatively homogeneous zircon with faint traces ofthe primary zoning (Fig. 10B), others are onlyaffected along the core-overgrowth boundary (Fig.10C). The BSE-bright replacements have distinctlyhigher [U] and lower Th/U ratios than the zonedcores; compare the marked spots in Fig. 10A cor-responding to n217-9a and 9b in Table 3. In con-trast to zircons in the less veined orthogneiss atGallared S (see above), the BSE-dark overgrowths


are often continuous around individual crystalsalthough variable in width. A second type of zirconis homogeneous or displays weak sector zoningand/or planar growth bands (Fig. 10D). This typeof zircon is relatively BSE dark, lacks overgrowthsas well as BSE-bright zircon replacements.

The oscillatory and sector zoned zircon coresare low in [U] and have Th/U ratios around 0.8(Fig. 7, Table 3). Four analyses of this type ofzircon yield an age of 1679�13 Ma (MSWD=0.2, Fig. 10E), interpreted to date the igneousprotolith of this vein sample. Three analyses gave

Fig. 10. (A) to (D) are back-scatter electron images of zircons from the vein sample at Gallared N. Abbreviations follow those inFig. 6. (A–C) show zircons variably replaced by BSE-bright zircon. Replacement is extensive in (A) and (B). In (A), BSE-brightzircon clearly truncates the primary zoning whereas in (B), the whole core is replaced by BSE-bright zircon. Note that oscillatoryzoning is discernible as ghost structures (arrows) in BSE-bright zircon. (C) This zircon show replacement essentially localised alongthe core-rim boundary. (D) diffuse (sector-) zoning in a homogeneous ca. 1.44 Ga zircon. Spot analyses in (A) are n217-09a and 09b,in (B) n217-06a and 06b and in (D) n217-03a (see Table 3). (E) ion microprobe data of oscillatory and sector zoned zircon, (F) ionmicroprobe data of homogenous zircon overgrowths, non-complex (homogeneous) zircons and BSE-bright replacements.


lower ages and were omitted on the basis of low206Pb/204Pb ratios and/or discordance (Table 3).

The similarities in U–Th composition, zoningtextures and age in samples from Gallared N andS, suggest that the protoliths may well be partof the same magmatic event. This interpretationis consistent with the similarities in rock com-position, field appearance and lack of anyobserved lithological discontinuity between theareas.

A concordant age of 1449�11 Ma (MSWD=10) is obtained by grouping U-Pb ages of, (1)BSE-dark overgrowths and homogeneous (non-complex) zircons and (2) distinct BSE-bright re-placements (Fig. 10F). Individual ages for thesegroups are 1452�7 (MSWD=0.4) and 1446�13 Ma (MSWD=12), respectively. Group (1) hasa wide range in Th/U ratios (Fig. 8); most analy-ses yield values higher than is typical for meta-morphic zircon. In contrast, the BSE-brightzircon replacements are consistently low in Th/U(�0.01) with [U] of 800–1600 ppm. Despite theinsignificant age difference between group (1) and(2) the marked differences in elemental (Th–U)composition imply that the zircons are not co-ge-netic. Textural relationships demonstrate thatgroup (1) is older (1452�7 Ma) than group(2)(1446�13 Ma). This means that the age differ-ence between the two zircon types is below theanalytical resolution. The age scatter (MSWD=12) for group (2) indicates formation of BSE-bright replacements at different stages orcontinuously over a significant time interval. Theorigin and U–Th signature of group (1) and (2)are discussed in more detail in 5.1 and 5.3 below.

4.1.2.2. Late-kinematic pegmatite. A pinch-and-swell shaped pegmatite, collected from a locallyderived large boulder, is bordered by a few mmthick biotite-rich margin to the host gneiss (Fig.9C). The pegmatite consists of microcline, quartz,biotite, minor homblende, opaques and abundantwhite plagioclase megacrysts, up to 5 cm large.With the exception of slightly undulose extinctionin quartz and feldspars, the minerals are un-strained. Late alterations include sericitization of

feldspar, growth of muscovite and replacement ofbiotite by chlorite and opaque.

Most zircons are prismatic, often with sharpedges. A majority contain cores (Fig. 11A and B)although homogeneous, non-complex, crystalsalso occur (Fig. 11C). The overgrowths and non-complex crystals vary from BSE-dark to -light,homogenous, except for a faint, irregular zoning.[U] ranges between ca. 100 and 1000 ppm whilethe Th/U ratio is consistently around 0.04 (Fig.7). Eight spot analyses yield an average 207Pb/206Pb age of 961�13 Ma (MSWD=2.6, Fig.11D), interpreted as dating crystallization of thislate-kinematic pegmatite and thus the final phaseof Sveconorwegian deformation in the Gallaredarea.

The inherited cores are generally replaced ex-tensively by BSE-bright, unzoned zircon similar tozircons in the vein sample. Analyses of such re-placements have 207Pb/206Pb ages of 1472�10,1436�30 and 1440�6 Ma. The first, slightlyolder, analysis is reversely discordant. Subsequentcheck in the electron microscope revealed that itwas acquired at the interface between BSE-brightand older (ca. 1.7 Ga) magmatic zircon and, thus,the age is meaningless.

4.2. Dagsas

Two samples of plagioclase-rich orthogneissfrom Dagsas have been investigated, one darkgrey, unveined gneiss and one with abundantreddish veins without melanosome (Fig. 12). Alsothe veins are strongly deformed; this is mostapparent in the veined gneiss where the veins aretransposed into complete parallelism with thegneissic foliation (Fig. 12). Mineral fabrics definea strong subhorizontal stretching lineation thattrends NE–ENE and a gneissic foliation that dipsonly gently to the SE or SSE.

4.2.1.1. Un�eined orthogneiss (SNG: 633315/130120). The sampled orthogneiss is homoge-neous and dark grey. In thin sections, thedeformation fabric is defined by trails and lensesof fine-grained aggregates rich in dark mineralsanastomosing with plagioclase-rich aggregates


Fig. 11. BSE-images of zircon in the late-kinematic pegmatite, Gallared N. Abbreviations follow those in Fig. 6. (A) and (B),prismatic overgrowth around xenocrystic cores that are extensively altered and/or replaced by BSE-bright zircon, (C) prismaticnon-complex zircon with sharp edges between crystal surfaces. Spot analyses are in (A) n218-09a and 09b, in (B) n218-l0a and in(C) n218-04a (Table 3). Note the varying BSE-colour of zircon formed during crystallization of the pegmatite, (D) U–Pb concordiadiagram of the late-kinematic pegmatite.

and ribbon- to lens-shaped quartz domains. Darkminerals are finegrained clinopyroxene, garnet,homblende, orthopyroxene and opaques. Quartzribbons are annealed; feldspars are granoblastic toirregular and antipertitic. Single, strained,megacrysts of non-perthitic plagioclase wrappedby the foliation may represent remnants of pri-mary igneous grains.

In contrast to zircons in the Gallared or-thogneisses, most grains have thick rims (Fig.13A). These are more or less homogeneous, BSE-light and vary in thickness between 20 and 40 �m.Magmatic cores are BSE-dark, homogeneous orhave oscillatory growth zoning (Fig. 13B). Theoriginal zoning in some crystals is altered byBSE-light spots often associated with radialcracks (Fig. 13A). Homogeneous, non-complex,BSE-light zircons also occur.

Six spot analyses of BSE-light overgrowths andnon-complex crystals plot concordantly at 1428�

9 Ma (MSWD= l.0, Fig. 13C). The low MSWDvalue suggests that overgrowths and non-complexcrystals were formed during the same event. De-spite the relatively high Th/U ratios around 0.5

Fig. 12. Plagioclase-rich orthogneiss with reddish veins atDagsas. Note the strong overprinting deformation fabric.


Fig. 13. U–Pb data and BSE-images of zircons from the unveined orthogneiss at Dagsas. Abbreviations follow those in Fig. 6. (A)thick BSE-light overgrowth around altered BSE-dark zircon (spot analyses n221-04a and 04b), (B) oscillatory zoned zircon (spotanalysis n221-03a), (C) ion microprobe data of rims and non-complex crystals, (D) concordia diagram showing conventional U–Pbdata.

(Fig. 14), often characteristic of magmatic growth,this generation of zircon is interpreted as havingformed during high-grade metamorphism, possi-bly in the presence of a magmatic fluid. Twoanalyses in the same grain yielded the youngest(n221-01a, Table 4) and the oldest (-01b) age ofall analyses; these were excluded from age calcula-tion. Ion probe data of oscillatory zoned zirconare treated together with analyses of similar zir-cons in the veined variety of the Dagsasorthogneiss.

Conventional U–Pb multi-grain analyses of zir-cons in the Dagsas orthogneiss sample yield inter-cept ages of 1711�48 and 1126�144 Ma(MSWD=3.7, Fig. 13D). The abraded fractionsare the least discordant (Table 1), by analogy tothe conventional dating of the Gallared or-thogneiss. As no Sveconorwegian-aged zircon isidentified in the ion probe data, mixing between ca.1.7 and 1.43 Ga zircon overprinted by a modernlead loss, is inferred to account for the poorlyconstrained lower intercept age of ca. 1.13 Ga.

4.2.1.2. Veined orthogneiss (SNG: 633280/130080).The plagioclase-rich gneiss is similar to the un-veined variety. The reddish veins are made up offeldspar domains (recrystallized aggregates withgranoblastic to anhedral grains) anastomosingwith very elongated ribbon quartz (annealed).

Fig. 14. [U] versus Th/U diagram for zircons from the plagio-clase othogneiss at Dagsas.

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Table 4Ion probe U–Pb data, Dagsas

Pb 207Pb/206Pb206Pb/204Pb 206Pb/238U 207Pb/235U D207Pb/206PbTh/U 206Pb/238UTh 207Pb/235UUGrain spot�1� (%) �1� (%) (%)age�2�(ppm) (Ma)�1� (%)(ppm) (Ma)(ppm)

Dagsas, un�eined orthogneiss

Oscillatory-zoned zircons87 20 200 0.2783 (3.4) 3.892 (3.5) 0.1014 (0.67) 1583 1612 1650�24 068 0.72n221-03a 32

58 28 600 0.3210(5.6) 4.573(5.6) 0.1033(64) 1795 1744 1685�24 −589n221-04b 0.701410.2951(3.3) 4.198 (3.4) 0.1032 (0.80) 1667 1674 1682�3040 600 025n221-05a* 64 47 0.720.2950 (2.9) 4.267 (3.0) 0.1049 (0.71) 1666 1687 1713�26n221-06a* 081 70 0.79 32 19 9000.2789 (3.6) 3.824 (3.7) 0.0995 (.76) 1586 1598 1614�28�105 0n221-08a 300.877579

71 �105 0.3041 (2.6) 4.378 (2.7) 0.1044 (0.63) 1711 1708 1704�24 065 0.87n221-10a* 29

O�ergrowths and homogeneous. non-complex zircons0.2426 (1.7) 2.906 (1.8) 0.0869 (0.45) 1400262 13830.54 1358�18 0505n221-01a 152 14 8000.2323 (1.7) 2.933 (1.8) 0.0916 (0.48) 1346n221-01b 1390233 1459�18 +5131 0.52 68 73 7000.2441 (1.7) 3.004 (1.8) 0.0893 (0.57) 1408 1409 1410�2297 400 0n221-02a* 1090.572083550.2414 (1.6) 2.986 (1.7) 0.0897 (0.52) 1394n221-04a* 1404317 1419�20 0164 0.49 95 75600.2476 (1.9) 3.090 (1.9) 0.0905 (35) 1426 1430 1436�14�105 0116n221-06b* 378 201 0.490.2419 (2.0) 3.007 (2.1) 0.0902 (0.47) 1396 1409 1429�18n221-07a* 0305 92 0.53 92 �105

0.2432 (4.0) 3.027 (4.0) 0.0902 (0.41) 1404 1414 1431�16�105 0n221-08b* 960.44153322124 59 700 0.2377 (1.6) 2.955 (1.6) 0.0902 (0.36) 1375 1396 1429�14 +1n221-09a* 423 219 0.47

Dagsas, �eined orthogneiss

Oscillatory-zoned cores0.2905 (1.5) 4.145 (1.6) 0.1035 (0.43)n219-03a* 1644161 1663 1688�16 0116 0.74 62 �105

0.2659 (1.9) 3.552 (2.0) 0.0969 (0.67)n219-05a 1520180 1539 1565�26 096 0.48 59 �105

0.2757 (2.7) 3.772 (2.7) 0.0993 (0.46) 1570 1587 1610�18�105 0n222-01a 630.951611640.2903 (2.1) 4.111 (2.2) 0.1027 (0.61) 1643n222-02a* 1656132 1673�22 0103 0.72 50 �105

0.2674 (3.1) 3.790 (3.2) 0.1028 (0.74) 1528 1591 1675�28�105 +431n222-05a* 79 101 1.10.3057 (3.2) 4.304 (3.3) 0.1021 (0.86) 1720 1694 1663�32 0n222-06a* 142 52 0.39 53 �105

O�ergrowths and non-complex ( fir-tree zoned) zircons0.2296 (2.4) 2.823 (2.5) 0.0892 (0.58) 133261 13620.35 1408�22 −1139n219-01a* 61 21 000

50 8800 0.2504 (2.2) 3.124 (2.3) 0.0905 (0.65) 1440 1439 1436�24 0n219-02a* 159138 1.20.2527 (4.2) 3.020 (4.3) 0.0867 (0.72) 1453 1413 1353�2840 600 0n219-04a 421.01071200.2653 (1.5) 3.090 (1.5) 0.0844 (2.0) 1517 1430 1303�80n222-01b −15453 56 0.04 136 12300.2395 (3.1) 2.895 (3.2) 0.0877 (0.60) 1384 1381 1375�2434 600 0n222-03a 641.8308164

n222-04a* 0.2329 (2.4)142 2.860 (2.5) 0.0891 (0.71) 1350 1371 1405�28 0179 1.1 48 47 9000.2394 (2.9) 2.972 (3.0) 0.0901 (0.79) 1383 1401 1427�3088 700 036n222-07a* 101 136 1.3

n231-01a 0.2167 (2.5)113 2.514 (2.6) 0.0842 (1.0) 1265 1276 1296�40 0180 1.5 38 4270



Fig. 15. U–Pb data and BSE-images of zircons from the vein-rich sample at Dagsas. Abbreviations follow those in Fig. 6. (A)oscillatory zoned zircon rimmed by a thick overgrowth (spot analyses n222-01a and 01b), (B) fir-tree zoned non-complex crystal(spot analysis n222-04a). Ion microprobe data of: (C) oscillatory sector zoned zircon, (D) overgrowths and non-complex zircon inthe vein-rich sample.

Most zircons are similar to those in the un-veined orthogneiss, and have BSE-light over-growths, although thinner (10–30 �m; Fig. 15A).A second type of zircon is non-complex and has aspecial type of sector zoning (Fig. 15B). Thesector boundaries form a zig-zag shape, calledfir-tree zoning (Raven and Dickson, 1989), andhave been suggested as reflecting fluctuation ingrowth rate during zircon crystallization (Vavra etal., 1996; their Fig. 2: 1–6).

Seven analyses of magmatically zoned zircon inboth unveined and veined orthogneiss samplesyield an age of 1686×14 Ma (MSWD= l.8, Fig.15C), interpreted as the protolith age of theDagsas orthogneiss. The zircons are consistentlylow in [U] with variable Th/U-ratios averaging 0.8(Fig. 14). Four analyses of BSE-light overgrowthsand fir-tree zoned single crystals result in an ageof 1419�12 Ma (MSWD= l.4, Fig. 15D). In-cluding analysis n222-03a which is analytically ofhigh quality (Table 4) yields a somewhat older ageof 1410�30 Ma and increases scatter (MSWD=

4). No obvious age differences persist betweenovergrowths and homogeneous zircon, nor be-tween the two Dagsas samples. The Th/U ratiosof the 1419�12 Ma zircons in the veined or-thogneiss are generally much higher than the1428�9 Ma zircons in the unveined orthogneisssample (Fig. 14). In the Dagsas area, metamor-phic reworking is thus dated at 1428�9 Mawhereas 1419�12 Ma is the preferred age fordating vein formation in the Dagsas area.

4.3. Mardakle�

The Mardaklev granitic gneiss is a fine-grained,in places aplite-like, sub-leucocratic rock thatforms an irregularly shaped unit in a several tensof km2 large area E and NE of Gallared (Fig. 3).Large portions of it may give the false impressionthat the rock virtually has escaped deformation.However, similar to the aplitic portions in thepre-tectonic dyke at Gallared S, this is due to thecombined effects of small grain size, leucocratic


composition (few dark minerals) and extensiverecrystallization.

A pinkish gneiss sample, with faint nebuliticvein structures, was collected in order to constrainthe protolith age and the event of reworking(SNG: 635040/132890). All zircons investigatedare complex and resemble those in the vein sam-ple at Gallared N, i.e. large portions of oscillatoryzoned zircon are replaced by homogeneous BSE-bright zircon that truncates the original magmaticzoning (Fig. 16A and B). Seven spot analyses ofoscillatory zoned zircon were excluded from cal-culation on the basis of low 206Pb/204Pb ratios(�8000) and/or being �5% discordant. The re-maining eight analyses of oscillatory zoned zircongive an age of 1676�10 Ma (MSWD=0.9, eightanalyses) whereas BSE-bright zircon give 1430�21 Ma (MSWD=5.2, calculated from four out ofseven analyses, Fig. 16C). These ages are inter-preted as dating the protolith of the Mardaklev

granitic gneiss and metamorphism, respectively.Three spot analyses of small, �20 �m, homoge-neous zircons yield Sveconorwegian ages. The low206Pb/204Pb recorded in these analyses (Table 5)probably reflects contamination by the epoxyresin. Some large zircons have thin, less than 10�m, BSE-light rims that may represent Sveconor-wegian overgrowths. They are, however, too thinto be dated.

The previous age estimate for this granitic body(1445�5 Ma, acquired from discordant, but ana-lytically precise, zircon multi-grain fractions;Welin, 1994) can thus be rejected as dating theprotolith. Both the ion microprobe result and theMSWD value of 350 in Welin’s data appear betterexplained by the zircon complexity noted here,with igneous ca. 1.68 Ga zircon, substantial por-tions of U- and Pb-rich, ca. 1.43 Ga, replacementsand a subordinate Sveconorwegian zirconcomponent.

Fig. 16. The Mardaklev gneiss sample. Abbreviations follow those in Fig. 6. (A) and (B) show extensive replacement of primaryzircon to BSE-bright zircon; spot analysis in (B) is n461-05a. (C) illustrates U–Pb ion microprobe data of complex zircons. Notethat a late Sveconorwegian zircon generation, in contrast to the Gallared and Dagsas samples, is detected in the Mardaklev gneisssample.

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Table 5Ion probe U–Pb data, Mardaklev gneissic granite * Analysis used in age calculation

PbTh 206Pb/204Pb 206Pb/238UUGrain spot 207Pb/235UTh/U 207Pb/206Pb D206Pb/238U 207Pb/206Pb207Pb/235Uage�2�(Ma) (%)�1� (%) (Ma)�1� (%)(ppm) (ppm) �1� (%)(ppm)

Mardakle�, gneiss sample

Oscillatory-zoned cores0.2624 (2.8) 3.745 (2.9) 0.1035 (1.0)n232-01a 1502117 1581 1688�38 +7123 0.89 44 11300.2179 (1.9) 2.819 (2.0) 0.0938 (0.78) 1271 1361 1505�303280 +1471n232-02a 277 74 0.200.2987 (2.2) 4.244 (2.3) 0.1031 (0.56) 1685 1683 1680�20n233-01a* 0190 174 0.87 77 11 0000.2717 (5.2) 3.889 (5.4) 0.1038 (1.5) 1549 1611 1693�56�105 0n234-01a* 161.35639

41 32 000 0.2741 (3.6) 3.914 (3.7) 0.1036 (0.90) 1561 1617 1689�34 +2n461-02a* 125105 1.00.2789 (2.3) 3.943 (2.8) 0.1025 (1.6) 1586 1623 1671�60880 +1n461-03a 291.080700.2852 (2.4) 3.956 (2.7) 0.1006 (1.3) 1617 1625 1635�46n461-04a 0149 197 1.3 65 17400.2860(0.71) 4.016 (0.95) 0.1018 (0.67) 1622 1637 1658�2577 000 +2n540-03a* 1010.6516271

31 000246 0.2802 (0.59) 3.970 (0.84) 0.1028 (0.59) 1595 1628 1675�22 +520 0.62 89n540-05a*0.2544 (0.66) 3.432 (1.3) 0.0979 (1.2) 1461 1512 1584�442630 +9207n540-06a 670.5615

68 000161 0.2822 (0.64) 4.011 (1.1) 0.1031 (0.91) 1602 1636 1681�33 +520 0.62 59n540-07a*0.2964 (0.61)n540-09a* 4.272 (1.3)137 0.1045 (1.1) 1673 1688 1706�40 +215 1.1 59 10 0000.2634 (0.59) 3.688 (2.0) 0.1016 (1.9) 1507 1569 1653�706700 +1032n591-02b 87 13 0.930.2826 (0.58) 3.979 (1.1) 0.1021 (0.91) 1605 1630 1662�33 +4n591-03a* 204 20 1.1 82 15 3000.2736 (0.61) 3.970 (1.1) 0.1052 (0.94) 1559 1628 1718�34 +1024 000n591-04a 101 12 1.1 41

Back-scatter bright zircon394 �105 0.2569 (1.2) 3.208 (1.2) 0.0906 (0.24) 1474 1459 1437�10 084n232-03a* 0.061385

0.2336 (0.91) 2.887 (0.95) 0.0896 (0.27) 1353 1378 1418�1017 100 +3953n461-01a* 2500.109386 000711 0.2398 (0.96) 2.877 (1.1) 0.0870 (0.42) 1386 1376 1360�16 066 0.09 190n461-05a

0.2399 (0.61) 3.008 (0.71) 0.0909 (0.36) 1386 1410 1445�14 +5n591-01a* 1347 14 0.09 361 53 0000.2313 (0.59) 2.858 (0.74) 0.0896 (0.44) 1341 1371 1418�1762 000 +5n591-02a* 3350.111212850.1808 (0.97) 2.051 (1.1) 0.0823 (0.62) 1072 1133 1252�24 +16n598-05a 1237 1928 0.08 249 62000.2148 (0.94) 2.631 (1.0) 0.0888 (0.40) 1254 1309 1401�15 +114200n598-06a 1101 784 0.10 266

Small homogeneous zircon**29101764 0.1555 (2.2) 1.520 (2.4) 0.0709 (1.2) 932 938 955�50 0149 0.081 305n233-02a

0.1616 (1.6) 1.551 (2.4) 0.0696 (4.4) 966 951 917�178230 −2246n233-03a 1116 397 0.31n598-01a 0.1721 (1.0)94 1.793 (1.5) 0.0756 (1.1) 1024 1043 1083�46 +686 1.1 19 26 600

**Only approximate ages are inferred.


Fig. 17. Histogram showing the distribution of all ion microprobe U–Pb analyses from Gallared, Dagsas and Mardaklev. Note thedistribution skewed towards lower ages for the ca. 1.7 (dark grey) and 1.46–1.42 (grey) Ga zircon populations. No analysis ofovergrowths, non-complex zircon or recrystallized zircon domain (grey) is older than ca. 1.49 Ga, leading to the interpretation thatthe 1.46–1.42 Ga event was the first affecting the ca. 1.7 Ga oscillatory and sector zoned zircons (dark grey). D, discordance.

5. Interpretation of different types of zircon byinternal morphology, Th–U chemistry and U–Pbisotopic data

Ion microprobe data on samples at Gallared N,Gallared S, Dagsas and Mardaklev define threeage groups which cluster at ca. 1.70–1.68, 1.46–1.42 and 0.96 Ga (Fig. 17). Although largelytreated as age groups below, it should be notedthat different samples within each group some-times have slight, but significant, age differences.

5.1. Formation of zircon during magmaticcrystallization

The 1.70–1.68 Ga zircons in the Gallared,Dagsas and Mardaklev orthogneisses have typi-cally magmatic Th/U ratios (�0.2) and oscilla-tory euhedral growth bands and/or sector zoning.Together with previously dated orthogneisses inthe Boras–Ulricehamn area (Connelly et al.,1996; Cornell et al., 1997) and in the Halmstadarea (Johansson 1998; Christoffel et al., 1999), the

investigated 1.70–1.68 Ga orthogneisses are theoldest intrusive rocks known in the Eastern Seg-ment south of Lake Vanern.

Similar, but more variable, Th/U ratios areobtained from the 1452�7 Ma overgrowths andhomogeneous zircons in the vein sample at Gal-lared N and the 1419�12 Ma overgrowths andfir-tree zoned zircons in the veined orthogneiss atDagsas (Figs. 7 and 14). Based on Th/U ratios�0.2, and field observations, these ages are inter-preted to date veining in the respective area.

5.2. New growth and recrystallization of zircon ina subsolidus metamorphic en�ironment

From petrographic studies, there are no signs ofpartial melting in the unveined orthogneiss sam-ples at Gallared S and Dagsas. Secondary zircon,dated at 1.46–1.42 Ga, is more abundant in theDagsas orthogneiss sample (represented by thickovergrowths and newly formed grains) than in theGallared sample. In the Dagsas sample, the Th/Uratios are relatively consistent (0.44–0.57)


whereas a much greater range is seen in theGallared sample (0.04–1.1) Altogether, this typeof 1.46–1.42 Ga zircon has Th/U ratios that liecloser to those typical for magmatic (�0.2)rather than for metamorphic zircon (Figs. 7 and14, Tables 3 and 4).

In the [U] versus Th/U diagram it is notewor-thy that secondary zircon formed in the unveinedorthogneiss sample at Gallared S (dated at1464�8 Ma) and the coeval, but magmatic, zir-cons in the vein sample at Gallared N togetherdefine a mixing curve with partly overlappingTh-U signatures (grey circles in Fig. 7); the sec-ondary zircon in the orthogneiss sample havingthe lowest Th/U ratios (�0.12, grey circlesmarked G in Fig. 7). Considering the abundanceof veins in the investigated areas and their in-ferred intrusive character, high-Th/U fluidsderived from dykes and veins may have pene-trated the wall rock. Mixing between magmatic(high-Th/U) and metamorphic (low-Th/U) fluids,resulting in local chemical variations during zir-con growth, would be compatible with the Th/Urange for secondary zircon in the gneissprotoliths.

5.3. Formation of back-scatter brightreplacements

The interpretation of BSE-bright zircon in thevein sample at Gallared N and the Mardaklevgneiss sample requires consideration of the fol-lowing observations, (1) BSE-bright zircon do-mains occur both in the magmatic cores and theovergrowths, the process leading to formation ofBSE-bright zircon must thus postdate the over-growths; (2) formation of BSE-bright zirconstarted preferentially at the core-overgrowthboundary and then migrated towards the centerof the crystal (Fig. 10 A and C); (3) transitionfrom primary zoned to BSE-bright zircon wasassociated with approximately one order of mag-nitude increase in [U]; (4) the process was efficientin both resetting the U–Pb system and in produc-ing stable zircon that escaped later isotopic distur-bance; (5) finally, and perhaps the most importantfeature, the primary igneous zoning is still dis-cernible within BSE-bright zircon domains.

Pidgeon (1992) described two type ofboundaries between unzoned and zoned zircon.The first type, referred to as a transgressiveboundary, is ‘where unzoned zircon forms discor-dant patches which appear to intrude the euhedralgrowth zones’. van Breemen et al. (1987) at-tributed similar type of structures to resorption ina melt followed by reprecipitation of new zircon.The second type, called transitional replacement,has been attributed to recrystallization and refersto progressive transition from zoned to unzonedzircon in which the faint traces of primary zoningare discernible (Black et al., 1986; Claoue-Long etal., 1988). Progressive recrystallization has beenobserved associated with exclusion of non-matrix(‘contaminant’, Pidgeon, 1992) elements and re-crystallized zircon is expected to have lower Th/Uratios as a consequence of the slightly higher ionicradius of Th relative U.

Textural observations (1–5 above) and the lowTh/U ratios recorded point to recrystallization,rather than resorption followed by reprecipita-tion, as the main process during formation ofBSE-bright replacements. However, the low Th/Uratios appear mainly controlled by input of Urather than preferential loss of Th over U; theseanalyses have relatively constant and low Th con-tents (Table 3). Extremely low Th/U composi-tions, such as those recorded here, are typical forsecondary zircon formed in granulite facies condi-tions and rarely found in igneous zircon, exceptfor zircon crystallized from late-stage partial melts(Zeitler and Chamberlain, 1991). A plausible sce-nario is that recrystallization took place in thepresence of a fluid strongly enriched in U.

5.4. Zircon growth during crystallization ofanatectic melts

The lack of zoning in the ca. 0.96 Ga non-com-plex zircons and thick overgrowths in the late-kinematic dyke at Gallared N, is interpreted toreflect kinetically fast and continuous growth in afluid-rich anatectic melt. This type of zircon isrecognized by high [U] and low Th/U ratios (�0.04). Similar low Th/U ratios (�0.09), charac-terize zircons in late Sveconorwegian leucosomesand pegmatoid fillings in boudinaged amphibo-


lites in the ca. 1.37 Ga Tjarnesjo granite (Fig. 3,Andersson et al., 1999) and the ca. 946 Ma peg-matite dyke at Steninge, near Halmstad (Fig. 1,Christoffel et al., 1999). Moller and Soderlund(1997) speculated that discordant ca. 955 Magranite-pegmatite dykes in the Gallared regionmay have originated by partial melting duringexhumation. That interpretation seems viable aszircon in the granite-pegmatite dykes also havelow Th/U ratios (�0.02, Moller and Soderlund,1997; Andersson et al., 1999). One explanation forthe consistently low Th/U composition recordedin zircon crystallized from late Sveconorwegiananatectic melts is that low-Th/U fluids, originat-ing during high-grade metamorphism of sur-rounding rocks, were entrapped in these melts.Friend et al. (1996) reported similar U contentsand Th/U ratios in zircon overgrowths and coresfrom a ca. 2.71 Ga diatexite formed by extensivemigmatisation of gneisses at Greenland. Accord-ing to the their results together with Th/U ratiosof Sveconorwegian zircon in SW Sweden, thedegree of partial melting is likely an importantfactor for the U–Th–Pb systematics in anatecticzircon given the strong partitioning of U and Thduring partial melting (see also Whitehouse et al.,2001).

The consistent and distinct difference in Th/Ubetween zircon crystallized from granitic 1.45–1.42 Ga veins and ca. 0.95 granite-pegmatitedykes is mainly attributed to different sources ofmagma production. An anatectic source for lateSveconorwegian granite-pegmatite dykes agreewith the low abundance of Sveconorwegian plu-tonic rocks in the Eastern Segment. Voluminous,ca. 1.45 Ga old, plutonic rocks have been dated inthe Blekinge region but have not yet been foundin the Eastern Segment as reworked equivalents.However, large granite intrusions may still exist atdepth and which acted as the source for thepresently exposed ca. 1.45 Ga dykes and veins.

6. The ability of zircon to record high-grademetamorphism

Many studies, focusing on the mechanisms offormation of secondary zircon, have attributed

new zircon growth to breakdown of Zr-bearingmagmatic phases such as ilmenite, pyroxene andamphibole (e.g. Stimac and Hickmott, 1994;Vavra et al., 1996; Scoates and Chamberlain,1997). Formation of secondary zircon has alsobeen attributed to breakdown of baddeleyite(ZrO2; Wingate et al., 1998). Considering theseminerals as the primary source of Zr, the greatestpotential to form secondary zircon should beduring the first metamorphic event that involvesbreakdown of the original mineralogy, since Zrincorporated into zircon during these reactions isthen difficult to remobilize. Later events may ormay not result in new zircon growth depending onthe supply of Zr from reacting minerals (Fraser etal., 1997), on the resorption and crystallization ofzircon during partial melting (Watson, 1996), oron the introduction of a fluid (Wayne and Sinha,1988; Ayers and Watson, 1991).

The near-penetrative gneissic deformation inthe Gallared area, which is part of the UDZ (Fig.3), was before this study bracketed between 972�14 Ma (Johansson et al., 2001), the age of meta-morphic zircon enclosed by prograde zonedgarnet in the eclogites, and ca. 955 Ma post-tec-tonic granite-pegmatite dykes (Moller and Soder-lund, 1997; Andersson et al., 1999). Thelate-kinematic pegmatite at Gallared N directlydates the deformation at 961�13 Ma. Furthersupport for a Sveconorwegian age of deformationin the Gallared area is the pre-tectonic style of the1.37 Ga Tjarnesjo granite (Andersson et al., 1999)and the 1444�8 Ma granite dyke at Gallared S(this study), which hosts 982�15 Ma metamor-phic zircon. In the Dagsas area, the deformationstructures and fabrics transposing the veins anddykes are of similar character as in the Gallaredarea; thus, a Sveconorwegian age of deformationis considered the most probable in the Dagsasarea as well.

Although compelling evidence for a high-gradeSveconorwegian imprint in the southern part ofthe Eastern Segment, which in the Gallared areaincluded metamorphism through eclogite-gran-ulite-upper amphibolite stages, secondary zirconyielding a Sveconorwegian age has not been ob-served in the 1.70–1.68 Ga Gallared and Dagsasorthogneisses. This is in marked contrast to the


abundance of Sveconorwegian secondary zircon inrocks younger than ca. 1.45 Ga in the EasternSegment (Fig. 2). Sveconorwegian growth of zirconessentially took place from late Sveconorwegiananatectic melts and in �1.45 Ga rocks metamor-phosed under subsolidus conditions-i.e. in the pre-tectonic ca. 1.44 Ga granitic dyke at Gallared S, inthe Varberg charnockite (L. Johansson, pers.comm.), and in many metabasic rocks (although ofunknown protolith ages, e.g. Wang et al., 1998;Johansson et al., 2001).

The distribution of secondary zircon in differentgenerations of intrusive rocks can be explained bya decreasing probability of metamorphic zircon toform as the rock undergoes repeated events ofmetamorphism. In the 1.70–1.68 Ga orthogneissesat Gallared and Dagsas, Zr hosted in igneousphases was liberated during the 1.46–1.42 Ga eventas the primary minerals recrystallized or took partin metamorphic reactions. The released Zr reactedwith silica to form new zircon grains as well asovergrowths around magmatic crystals. Thus, inthe older rocks, there was less easily mobilised Zrleft for a new generation of secondary zircon toform during the subsequent Sveconorwegian event.It is possible that extensive partial melting (or fluidinfiltration) of the oldest, already metamorphosed,rocks was necessary to liberate sufficient amountsof Zr for new zircon to form during the Sveconor-wegian. This conforms with the occasional presenceof Sveconorwegian zircon in rocks younger than ca.1.45 Ga and in rocks generated from anatecticmelting (e.g. the dated ca. 961 Ma late-kinematicpegmatite and slightly younger ca. 955 Ma granite-pegmatite dykes in the Gallared area). To ourknowledge, this is one of the first well-documentedcase where secondary zircon did not form despitecomplete recrystallization, penetrative deformationand metamorphism under high-grade, near-anatec-tic, conditions.

7. Regional implications

7.1. Pre-S�econorwegian reworking at 1.46–1.42Ga: a static or dynamic thermal e�ent?

Hubbard (1975) assigned high-grade metamor-

phism, anatexis and emplacement of granites,charnockites and mafic rocks in the Varberg regionto an orogenic event at ca. 1.4 Ga which he namedthe ‘Hallandian event’. However, in the Hallandianmodel the effects of Sveconorwegian high-pressuremetamorphism, anatexis and deformation were notconsidered. From work in the Varberg–Halmstadregion, Christoffel et al. (1999) re-assessed the termHallandian to denote a thermo-magmatic event inthe 1.44–1.38 Ga interval.

In accordance with Christoffel et al. (1999), theterm ‘Hallandian’ is maintained by us to describethermo-magmatic features in the Gallared, Dagsasand Mardaklev areas, and without implying anecessary linkage to an orogenic cycle as proposedby Hubbard (1975). Hallandian activity appears tohave been slightly younger in the Dagsas area(1419�12 and 1428�9 Ma) than in the Gallaredarea (1446�13 and 1464�8 Ma). The lower agelimit of ca. 1.38 Ga for Hallandian activity in theVarberg–Halmstad, reported by Christoffel et al.(1999), refers to crystallisation ages of the Torpaand Varberg intrusives. It is possible that these1.38–1.37 Ga intrusive rocks represent the finalpulse in a prolonged ca. 80 Ma magmatic period,alternatively they belong to a separate, slightlyyounger magmatic event. Here, the preferred agelimits for Hallandian thermo-magmatic activityrefer to the ages of secondary zircon growth andrecrystallization. If correct, the age of the mafic-fel-sic Axamo dykes (1410�10 Ma, Lundqvist, 1996)near Vrana (Fig. 1) may represent the end ofHallandian activity.

Emplacement of dykes and formation of sec-ondary zircon at 1.46–1.40 Ga have been recog-nized from Halmstad in the SW to Vrana in theNE (Fig. 1, Soderlund, 1996; Connelly et al.,1996; Johansson, 1998; Cornell et al., 1999;Christoffel et al., 1999), coeval with 1.45 Gagranite magmatism in SE Sweden (Kornfalt andVaasjoki, 1999, and references therein). From atectonic perspective it is a critical questionwhether or not the 1.46–1.42 Ga event was asso-ciated with regional penetrative deformation. Pre-Sveconorwegian deformation has been found inthe Halmstad area and at Vrana (Fig. 1), repre-sented by ca. 1.46–1.40 Ga, variably deformed


dykes that cut across a deformational, generallymigmatitic, fabric (Connelly et al., 1996; Christ-offel et al., 1999). Pre-Sveconorwegian deforma-tion is also apparent at Gallared as the ca. 1.44Ga deformed granite dyke, in which surfaces ex-posed at high angles to the lineation, reveal apreserved low angle discordant relationship (Fig.4A). Thus, these localities record early Hallan-dian deformation and migmatisation or stillolder events, see below.

The 1.46–1.42 Ga event in SW Sweden is co-eval with granulite facies metamorphism in theSW Grenville Province (Ketchum et al., 1994),and slightly younger than 1.50–1.46 Ga tecton-ism in the Grenville Province (‘Pinwarianorogeny’, Rivers, 1997; Kamo et al., 1996).From a circum-Atlantic tectonic and correlativeperspective, the 1.46–1.42 Ga event clearly mer-its further attention. Further characterization ofthis event requires investigation of areas thathave escaped penetrative Sveconorwegian defor-mation. The best opportunities are in easternparts of the Eastern Segment (in areas close tothe SFDZ), where Sveconorwegian reworkingwas less penetrative, or, even better, eastthereof.

7.2. Implications of zircon genesis for constrainingthe tectono-metamorphic e�olution of the EasternSegment

Christoffel et al. (1999) reported a zircon ageof 1654�9 Ma from a boudinaged and amphi-bolitized mafic dyke at Steninge situated in amigmatitic orthogneiss. Based on structural in-terpretations (e.g. the lack of leucosomes in themetabasite) the mafic dyking was interpreted aspostdating an event of deformation and migma-tization in that area. The age of metamorphiczircon (1659�5 Ma) is identical within errorwith the age of the gneiss protolith (1664�7Ma) and Christoffel et al. (1999) reviewed alter-native interpretations. Based on cathodolumines-cense images, the zircons were interpreted to bemetamorphic and directly date Gothian rework-ing at 1659�5 Ma. At present, this is the onlydate of secondary zircon in the Eastern Segmentwhich yields a Gothian age. Considering the

possibility (or probability) of a very subtle agedifference between discordant ca. 1.46–1.40 Gadykes (at Vrana, Steninge, and possibly also atGallared S) and deformation fabrics in the hostrocks, there is very little support for a Gothian(�1.55 Ga) gneiss-forming event in the EasternSegment.

Tectonic models have attributed the westwardyounging of magmatism to discrete events ofoceanward stepping subduction during the 1.75–1.55 Ga Gothian orogeny (A� hall and Gower,1997; A� hall and Larson, 2000). Subduction re-sulted in successive accretion of arc systems (i.e.Horred and SLM belts of the Idefjorden ter-rane), onto a proto-margin of Baltica (EasternSegment) prior to ca. 1.62 and 1.59 Ga, respec-tively. During these processes substantial re-working took place in rocks east of thecollisional boundary, namely in the 1.70–1.68Ga orthogneisses investigated here (e.g. A� halland Gower, 1997; Brewer et al., 1998; A� hall andLarson, 2000).

These tectonic models are not compatible withpresent geochronological data of deformationand metamorphism in the Eastern Segment(Figs. 2 and 17), and certainly not with themechanisms controlling zircon growth as dis-cussed here. If they were correct, new zirconyielding Gothian ages should have formed, andprobably in significant amounts. It is unlikelythat Gothian zircon rims and newly formed zir-con grains were completely and selectively dis-solved during the succeeding 1.46–1.42 Ga andSveconorwegian events. Although zircon coressometime show signs of resorption or recrystal-lization, these processes are generally restrictedto parts of a crystal, and many igneous zirconsin the orthogneisses have well-preserved pris-matic shapes, in disagreement with substantialpost-magmatic resorption. Complete resettingand resorption may occur locally, e.g. in thepresence of acid fluids or in rock volumes thatundergo extensive melting (Ayers and Watson,1991), but not on a regional scale resettingwhole zircon populations in rocks of differentcompositions.


8. Concluding remarks

We have dated complex zircons in variablyreworked and veined orthogneisses, representingthe oldest known lithologies in the southern, high-grade part of the Eastern Segment, SW Sweden.Magmatically zoned zircon yields ages of 1.70–1.68 Ga and date the protoliths to the or-thogneisses. Secondary zircon in orthogneisssamples and zircon grown during veining yieldsages of 1.46–1.42 Ga. This event largely coincidesin age with the Hallandian event, originallydefined as a complex orogenic cycle in the Var-berg region (Hubbard, 1975). The term ‘Hallan-dian’ is by us restricted to describe featuresassociated with 1.46–1.42 Ga thermo-magmaticactivity, and not Sveconrwegian, in the studyareas. Consistent transposition of 1.46–1.42 Gadykes and veins, and reworking of ca. 1.37 Gaplutonic rocks indicate Sveconorwegian deforma-tion which, in one area (Gallared), includes well-documented metamorphism through eclogite-granulite-upper amphibolite stages. Prior to thisstudy the Sveconorwegian deformation in theGallared area was bracketed between ca. 972 and955 Ma and the final phase is here directly datedat 961�13 Ma.

The lack of secondary Sveconorwegian zirconin the 1.70–1.68 Ga Gallared and Dagsas or-thogneisses, is consistent with igneous mineralsbeing the most important source of Zr for theformation of new zircon under subsolidus meta-morphism. Zr hosted as a trace element in igneousminerals was redistributed into new zircon duringthe 1.46–1.42 Ga event. Thus, there was lesseasily mobilised Zr available for a new generationof secondary zircon to form during the Sveconor-wegian. We conclude that new zircon growthduring penetrative metamorphism and deforma-tion (even close to near-anatectic conditions) maybe inhibited in rocks that have experienced anearlier event of metamorphism.

Previous tectonic models which attribute crustalgrowth and reworking of the Eastern Segment toGothian (1.75–1.55 Ga), accretional-related oro-genic activity are not compatible by our data.Deformational and migmatite structures cross-cutby 1.46–1.40 Ga felsic dykes, as observed in a few

places in the Eastern Segment, have previouslybeen interpreted in favour of a major Gothiangneiss-forming event. In view of the present data,these structural-intrusive relationships may bebetter explained by dynamic reworking at ca. 1.45Ga.

Acknowledgements

Valuable improvements of the manuscript bythe journal reviewers K.-I. A� hall and B. Bingenare gratefully acknowledged. Comments by S.Bogdanova was also helpful. We thank T. Sundeand J. Vestin (NORDSIM, Stockholm) for techni-cal assistance at the ion microprobe. Financialsupport for this study was provided by theSwedish Natural Science Research Council(NFR), through grants to C. Moller, L. Jo-hansson and U. Soderlund. NORDSIM contribu-tion number 46.

Appendix A

Ion microprobe analyses.The U–Pb dating of the zircons was performed

using a CAMECA IMS 1270 ion probe at theNordsim laboratory at the Museum of NaturalHistory in Stockholm. Analytical procedures es-sentially followed those described by Schumacheret al. (1994) and Whitehouse et al. (1997) White-house et al. (1999). Errors in 207Pb/206Pb ratios areobserved analytical errors which generally exceedcounting statistic errors. Errors in Pb/U ratiosinclude a dominant component propagated fromreproducibility of the standard measurements. Th/U ratios have been estimated both from measured206Pb/208Pb ratios, to calculate a Th/U ratio for anassumed single stage evolution from the 207Pb/206Pb age, and from measured Th/U ratios, cali-brated to standard, yielding the present-day (i.e.measured) Th/U ratio.

Con�entional U–Pb analyses.The chemical procedures and mass spectrome-

try analysis were carried out at the Laboratory forisotope geology at the Museum of Natural His-


tory in Stockholm. Prior to decomposition, allfractions were washed in distilled acetone for 10min in an ultrasonic bath, and in weak HNO3

for 15 min and finally rinsed several times withdouble distilled water. The zircons were dis-solved in teflon bombs at 210 °C in a 10:1 mix-ture of HF and HNO3 for 5 days.

After aliquotation, a mixed 233–235U–208Pbspike was added to the isotope dilution aliquots.The ion-exchange was carried out using HCltechnique. The isotope composition and isotopedilution parts were dissolved i 2.0 and 3.1 NHCl, respectively, before the samples wereloaded in 50 �l columns. Total blank is esti-mated to 10 pg.

U and Pb were loaded on double and singleRe-filaments, respectively (Pb together with sil-ica gel and phosphoric acid), and analyzed usinga Finnigan MAT 261 mass spectrometer. Datareduction followed Ludwig (1980) using programversions by Ludwig (1991a) Ludwig (1991b) andthe decay constants given in Steiger and Jager(1977), except for samples in which a majorityof analyses plot concordant; for these, ages werecalculated as weighted means of individual207Pb/206Pb ages. The common Pb correctionswere based on the Pb isotopic evolution modelof Stacey and Kramers (1975). All presentedages are given at 2� confidence level.

Pb–Pb e�aporation analysis.The single-zircon evaporation analyses were

performed on the same mass spectrometer asused for conventional U–Pb analysis. Data wasacquired by magnetic peak switching of themass sequence 206–207–208–206–204 using asecondary electron multiplier. No correction wasmade for mass fractionation. Correction for ini-tial lead was based on measured 206Pb/204Pb ra-tio for each data block with initial lead isotoperatios following Stacey and Kramers (1975). Zir-con ages were calculated as arithmetic means ofevaporation steps considered to represent analy-sis of undisturbed zircon (i.e. evaporation stepswithout any significant change in 207Pb/206Pbage).

Procedures of data selection-significance of asym-metric age distributions.

From Fig. 17, it is apparent that the twoolder age groups, dated at ca. 1.70–1.68 and1.46–1.42 Ga, do not follow a normal Gaussiandistribution but are skewed towards lower ages.Similar patterns have been recognized in ion mi-croprobe data from high-grade metamorphic re-gions elsewhere (e.g. Whitehouse et al., 1997).The skewed distribution has commonly been at-tributed to ancient Pb-loss, and analyses yieldingyoung ages may in this case be rejected in theage calculation.

In this study, analyses with 206Pb/204Pb ratiosbelow 8000 and/or those that are more than are�5% discordant, have been omitted from agecalculation as discordant data and low 206Pb/204Pb ratios generally indicate isotopic distur-bance. The chosen values permit the majority ofanalysis giving ‘too young’ ages to be excluded.Despite this, some concordant to nearly concor-dant analyses with high 206Pb/204Pb ratios re-main yielding ages significantly younger relativethe common mean of a specific zircon popula-tion in a sample. These analyses are here con-sidered justified to omit from age calculation as,(1) All the zircons have been carefully examinedusing back-scatter imaging. Analyses that yieldsignificantly lower ages than their commonmean in the 1.70–1.68 Ga populations are with-out any exception analyses of oscillatory or sec-tor zoned zircon; thus, they cannot be attributedto record a younger event of metamorphiczircon growth. (2) Analyses yielding ‘too young’ages do not cluster at any specific age thatcould correspond to a discrete geological event.(3) Since the precision in measured Pb/U ages isrelatively poor, a specific analysis with anerror ellipse overlapping the concordia trajectorymust not necessarily represent analysis of azircon volume that escaped isotopic disturbance,i.e. is ‘truly’ concordant. (4) There are twowell documented events of reworking in thestudy areas (at 1.46–1.42 Ga and the Sve-conorwegian) that potentially could have causedPb-loss.


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