Hallandian 1.45 Ga high-temperature metamorphism in Baltica: P–T evolution and SIMS U–Pb zircon ages of aluminous gneisses, SW Sweden

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Hallandian 1.45 Ga high-temperature metamorphism in Baltica: P–T evolution and SIMS U–Pb zircon ages of aluminous gneisses, SW Sweden Jan Ulmius a,∗ , Jenny Andersson b , Charlotte Möller a a b

Department of Geology, Lund University, Sölvegatan 12, SE-223 62 Lund, Sweden Geological Survey of Sweden, Box 670, SE-751 28 Uppsala, Sweden

a r t i c l e

i n f o

Article history: Received 4 September 2014 Received in revised form 26 March 2015 Accepted 1 April 2015 Available online xxx Keywords: Mesoproterozoic Accretionary Baltica Hallandian Polymetamorphism High-temperature metamorphism

a b s t r a c t The southernmost Baltic Shield exposes polymetamorphic continental crust that was largely formed and accreted during a series of 1.92–1.66 Ga Paleoproterozoic orogenic events and later reworked during the 1.14–0.90 Ga Sveconorwegian orogeny. An intermediate period of metamorphism, deformation and magmatism at 1.47–1.38 Ga has been attributed to the Hallandian orogeny, but due to overprinting by Sveconorwegian high-grade metamorphism and deformation, the P–T-t evolution and deformation of the Hallandian event have remained obscure. This study presents the first quantitative P–T model of the Hallandian event using high-temperature aluminous gneisses in the south-easternmost marginal part of the Sveconorwegian orogen. The high-grade metamorphism and spatially associated granite magmatism are dated using U–Pb SIMS analysis of zircon. Petrography, bulk and mineral geochemistry, and pseudosection models demonstrate prograde staurolite-sillimanite-grade metamorphism reaching granulite-facies temperatures (700–750 ◦ C) at low pressures (4–5 kbar), with the formation of Crd + Sil + Grt + K-fsp + Ilm + Melt ± Bt. The rocks followed a clockwise P–T path. Later stages involved the formation of sillimanite + biotite at the expense of garnet and cordierite. Local low-temperature and fluidassisted retrogression also caused the formation of chlorite and muscovite at the expense of cordierite. Both granite and aluminous gneisses contain complex zircon with inherited 1.70 Ga igneous cores and high-U, secondary zircon, mainly formed by reworking of protolith cores. The latter date the Hallandian high-grade metamorphism at 1451 ± 6 Ma and the granite magmatism at 1445 ± 8 Ma. The presence of 1.70 Ga igneous zircon cores in both metamorphic and magmatic rocks suggests that they formed from similar protoliths. The protolith ages correlate with the youngest generation of magmatic rocks of the Transscandinavian Igneous Belt. The aluminous gneisses are of supracrustal origin, and may have formed by chemical alteration of magmatic rocks. Hallandian regional metamorphism took place under a strongly elevated geotherm and was associated with granitic magmatism, suggesting an accretionary orogenic setting. The Hallandian event may demonstrate an 1.47–1.38 Ga Andean-type continental margin at the SW margin of Baltica. © 2015 Elsevier B.V. All rights reserved.

1. Introduction Accurate resolution of different metamorphic and magmatic components in Precambrian polymetamorphic shield areas is a challenge that needs to be met to understand evolution of continents (e.g., Hand et al., 1992, 1994; Engvik et al., 2000; Clark and Hand, 2010). The southwest Baltic Shield is an example of a large tract of composite polymetamorphic continental crust,

∗ Corresponding author. Tel.: +46 709490212. E-mail address: [email protected] (J. Ulmius).

which received its youngest metamorphic imprint during the 1.14–0.90 Ga Sveconorwegian orogeny (Bingen et al., 2008a). Large volumes of crust first formed and accreted during the Paleoproterozoic at 2.0–1.7 Ga (e.g., Stephens et al., 2009), and several later periods of metamorphism and magmatism between 1.6 and 1.2 Ga are recognised in different Sveconorwegian terranes (Bingen et al., 2008b, and references therein). In the case of the Baltic Shield, the introduction of SIMS spot dating of zircon (Nordsim laboratory, late 1990’s), with its capability of pin-pointing single stages of igneous and metamorphic zircon growth, meant a huge step forward. For example, the Eastern Segment of the Sveconorwegian orogen (Fig. 1) changed status

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Please cite this article in press as: Ulmius, J., et al., Hallandian 1.45 Ga high-temperature metamorphism in Baltica: P–T evolution and SIMS U–Pb zircon ages of aluminous gneisses, SW Sweden. Precambrian Res. (2015), http://dx.doi.org/10.1016/j.precamres.2015.04.004

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Fig. 1. Schematic overview of relict orogenic belts in the southern and central Baltic Shield including outline of 1.7–1.5 Ga rapakivi complexes and unmetamorphosed Mesoproterozoic igneous and sedimentary rocks in the central shield area. Figure based on 1:5 M Fennoscandian map database, and Geological Survey of Sweden 1:1 M Bedrock map database. Outline of the Oslo graben, the Sorgenfrei-Tornquist Zone and the southern Caledonian Deformation Front based on Erlström et al. (2004). Inset of Baltica from Bingen et al. (2008b).

Please cite this article in press as: Ulmius, J., et al., Hallandian 1.45 Ga high-temperature metamorphism in Baltica: P–T evolution and SIMS U–Pb zircon ages of aluminous gneisses, SW Sweden. Precambrian Res. (2015), http://dx.doi.org/10.1016/j.precamres.2015.04.004

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Fig. 2. Precambrian provinces in the southernmost Baltic Shield showing larger bodies of 1.47–1.38 Ga intrusions and metamorphic equivalents, compiled after the 1:5 M scale Fennoscandian Map data base and SGU bed rock map databases.

from a “Gothian” 1.75–1.55 Ga metamorphic domain locally overprinted by Sveconorwegian shear zones (cp. Connelly et al., 1996, and references therein) to a Sveconorwegian high-grade metamorphic domain with an older, cryptic, Hallandian metamorphic history at 1.47–1.38 Ga (Johansson et al., 2001; Andersson et al., 2002; Söderlund et al., 2002). Despite the advancements made possible by spot dating techniques, challenges remain in many cases to provide a firm petrologic link between age and metamorphic or magmatic stage, rendering the meaning of the age enigmatic. In this paper we present detailed petrology and P–T pseudosection models of pre-Sveconorwegian high-temperature aluminous gneisses in the southeasternmost marginal part of the Sveconorwegian orogen. These gneisses escaped reworking during the Sveconorwegian orogeny, and were only overprinted by nonpenetrative Sveconorwegian greenschist-facies metamorphism. The high-grade metamorphism and the spatially associated granite magmatism are dated using U–Pb SIMS analysis of zircon. This data provide the first well-characterised example of 1.45 Ga Hallandian metamorphism in the southernmost Baltic Shield, demonstrating it as an orogenic event with a clockwise P–T-t evolution in a hightemperature, low-intermediate pressure regime associated with granitic magmatism. The geographical extent and character of the Hallandian high-grade metamorphism, anatexis and magmatism is discussed, and an Andean-type continental margin is suggested as the plausible tectonic context. Together with other distinguishing features, the Hallandian 1.47–1.38 Ga orogenic imprint is absent in western Sveconorwegian terranes, emphasising the role of the

Mylonite Zone as a first-order lithotectonic collisional boundary of Sveconorwegian age. 2. Geological setting 2.1. Orogenic provinces in the southernmost Baltic Shield The southernmost Baltic Shield (Fig. 1) is made up of the eroded remnants of three Precambrian orogenic belts: a 1.92–1.66 Ga Paleoproterozoic orogen, the 1.47–1.38 Ga Hallandian orogen, and the 1.14–0.90 Ga Sveconorwegian orogen. Each of them are incorporated in, and partly overprinted by, younger orogenic systems (Hallandian, Sveconorwegian, and/or Caledonian; Fig. 1). The boundaries between the different orogenic provinces are defined by the limit of ductile deformation associated with the youngest orogenic event. In particular the Hallandian orogen has been almost completely reworked by a younger high-grade Sveconorwegian event, leaving the Blekinge-Bornholm province (Fig. 2) as its only preserved exposure. 2.1.1. The 1.92–1.66 Ga Paleoproterozoic province The 1.92–1.66 Ga Paleoproterozoic province in the southernmost Baltic Shield (Fig. 2) exposes bedrock formed and variably metamorphosed and deformed along a long-lived active continental margin (Mansfeld, 1996; Stephens et al., 2009; Bogdanova et al., 2015; Kleinhanns et al., 2015). It includes Paleoproterozoic rocks reworked and/or formed in a series of accretionary orogenic events at 1.92–1.79 Ga (Svecofennian orogeny; Lahtinen

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et al., 2005), and wide-spread alkali-calcic, granitic to quartzmonzodioritic magmatic rocks (Transscandinavian Igneous Belt, stippled in Figs. 1 and 2) formed along the same continental margin of Baltica at 1.87–1.85 Ga, 1.81–1.77 Ga and 1.74–1.66 Ga, broadly younging towards the southwest, and interpreted as formed at an Andean-type margin (Andersson et al., 2007; Andersen et al., 2009; Kleinhanns et al., 2015). The Paleoproterozoic metamorphic imprint in the southern shield area (Fig. 2) varies but is characterised by low-pressure conditions at around 5 kbar (Stephens et al., 2009). It is generally characterised by low-grade (greenschist-facies) conditions, although high-grade metamorphism with partial melting (T ≥630 ◦ C), was reached in some supracrustal complexes (Beunk and Page, 2001). 2.1.2. The 1.47–1.40 Ga Blekinge-Bornholm province The Blekinge-Bornholm province (Fig. 2) is made up of Paleoproterozoic 1.82–1.74 and Mesoproterozoic 1.47–1.44 Ga meta-intrusive rocks and gneisses, the older group also including meta-volcanic rocks (the Västanå association) and other gneisses of presumed supracrustal origin (Blekinge coastal gneisses; Fig. 2; Zarin¸sˇ and Johansson, 2009; Johansson et al., 2006, 2015). The 1.47–1.44 Ga igneous rocks are dominantly granitic-quartzmonzodioritic intrusions, which are partly well preserved (e.g., ˇ cys ˇ cys et al., 2002; Ceˇ Eringsboda and Karlshamn intrusions; cf. Ceˇ and Benn, 2007). Equivalents of these are also present in the subsurface East European Craton (e.g., Obst et al., 2004; Motuza et al., 2006). The rocks of the Blekinge Bornholm province was heterogeˇ cys et al., 2002; Johansson neously deformed at 1.47–1.44 Ga (Ceˇ ˇ cys and Benn, 2007; Zarin¸sˇ and Johansson, 2009). This et al., 2006; Ceˇ tectonometamorphic and magmatic event is part of the regionally extensive high-grade metamorphic event originally attributed to the 1.47–1.38 Ga Hallandian orogeny (Hubbard, 1975; Söderlund et al., 2002; Möller et al., 2007; Brander et al., 2012). Pervasively gneissic, in places migmatitic, varieties of the 1.47–1.44 Ga intrusions occur at the island of Bornholm (Zarin¸sˇ and Johansson, 2009; Waight et al., 2012). A close temporal relation between metamorphism and magmatism in the Blekinge-Bornholm province is indicated by 40 Ar–39 Ar cooling ages of hornblende and biotite at 1.45 Ga and 1.40 Ga, respectively (Bogdanova, 2001; Waight et al., 2012). Similarities between the orientation of structures in the 1.82–1.74 Ga metamorphic country rocks and anisotropy of the magnetic susceptibility inside the 1.47–1.44 Ga intrusions provide support for syn-deformational emplacement of the younger intruˇ cys and Benn, 2007). sions (Ceˇ In this paper we use the term Hallandian consistently for 1.47–1.38 Ga metamorphism, deformation and magmatism in the Baltic Shield and sub-surface counterparts in the East European Craton. This term was first established in the literature to denote 1.5–1.4 Ga magmatism and accompanying metamorphism in Scandinavia [cf. e.g., Hubbard (1975), and Gaál and Gorbatschev (1987)], and we find its original definition consistent with the present data. Bogdanova (2001) introduced the term “Danopolonian” to include, into the 1.5–1.4 orogenic context, 1.54–1.50 Ga anorthosite, mangerite, charnockite and granite intrusions in the sub-surface East European Craton; a term later used to denote also 1.47–1.38 Ga old metamorphism and magmatism synonymous with the already ˇ cys and Benn, 2007; Brander and established term Hallandian (Ceˇ Söderlund, 2009). 2.1.3. The 1.14–0.90 Ga Sveconorwegian province In late Mesoproterozoic to early Neoproterozoic time, Baltica was modified by the 1.14–0.90 Sveconorwegian collisional orogeny (Figs. 1 and 2; Berthelsen, 1980; Bingen et al., 2008a,b; Möller et al., 2015). Today the Sveconorwegian province forms a c. 500 km wide orogenic belt extending across southwestern Scandinavia,

the remainder of a 1.14–0.90 Ga large hot orogen. A conspicuous late Sveconorwegian deformation belt, the Mylonite Zone, defines a major lithological, structural and metamorphic terrane boundary dividing the Sveconorwegian province in an eastern and a western part (Fig. 1; Stephens et al., 1996; Andersson et al., 2002; Bingen et al., 2008a,b; Viola et al., 2011). The Eastern Segment is composed of reworked 1.92–1.66 Ga Palaeoproterozoic and 1.47–1.38 Ga Mesoproterozoic crust, continuous with lithologies east of the Sveconorwegian Front (Fig. 2; Söderlund et al., 1999; Petersson et al., 2013). The western Sveconorwegian terranes (Fig. 1) are composed of bedrock principally formed along 1.64–1.52 Ga Gothian (Bingen et al., 2005; Åhäll and Connelly, 2008) and 1.52–1.47 Ga Telemarkian (Bingen et al., 2008a) active margins. Their pre-Sveconorwegian locations and relation to the pre-Sveconorwegian Baltica is unclear (Andersson et al., 2002; Möller et al., 2015). With the exception of late-orogenic dykes, Sveconorwegian magmatism was limited to western parts of the orogen, while the eastern part of the orogen experienced extensive tectonic reworking, including deep burial and eclogitisation of Baltica continental crust. The Eastern Segment of the Sveconorwegian province is semi-continuous with the foreland to the east, and is almost entirely composed of deformed and metamorphosed equivalents of 1.87–1.66 Ga Transscandinavian Igneous Belt rocks (Fig. 1; Söderlund et al., 1999; Möller et al., 2007; Stephens et al., 2009; Petersson et al., 2013). Eclogite-facies metamorphism of parts of the Eastern Segment attests to deep tectonic burial (>55 km) of continental crust at 0.99–0.98 Ga (Möller, 1998, 1999; Johansson et al., 2001; Austin Hegardt et al., 2005; Möller et al., 2015). It was followed by high-grade metamorphism, ductile deformation and partial melting at 0.98–0.96 Ga (op. cit; Andersson et al., 1999, 2002; Söderlund et al., 2002; Möller et al., 2007). The Sveconorwegian high-pressure and high-temperature metamorphic reworking caused pervasive transposition and metamorphic overprinting of pre-existing migmatitic structures and assemblages, and resetting of most isotopic systems, including U–Pb in titanite (Wang et al., 1998; Söderlund et al., 2002; Möller et al., 2007; Johansson et al., 2013). The robust U–Pb system in zircon, however, preserves records of Hallandian high-grade metamorphism, in particular partial melting, also within the high-grade parts of the Eastern Segment (Table 1, Fig. 2; Christoffel et al., 1999; Söderlund et al., 2002; Möller et al., 2007; Brander et al., 2012). The late-Hallandian evolution recorded in the Eastern Segment demonstrates that the Hallandian orogeny terminated with granite-norite-charnockite magmatism, locally associated with fluid-assisted metamorphism, at 1.42–1.38 Ga (Hubbard, 1975; Åhäll et al., 1997; Christoffel et al., 1999; Andersson et al., 1999; Rimˇsa et al., 2007; Harlov et al., 2006, 2013; Johansson et al., 2013). The Eastern Segment thus comprises polymetamorphic highgrade metamorphic complexes that record extensive high-grade metamorphism at both 1.47–1.38 Ga and 0.99–0.96 Ga. Sveconorwegian metamorphic temperatures and deformational overprint decrease towards the east, and ductile deformation zones become successively more spaced towards the Sveconorwegian Front. Well preserved pre-Sveconorwegian high-grade metamorphic complexes occur along the Sveconorwegian Front, where Sveconorwegian overprint is limited to non-penetrative ductile deformation and metamorphism in lower amphibolite- or greenschist-facies; one such area is the Romeleåsen horst (Figs. 2 and 3). Metamorphic complexes along the Sveconorwegian Front and east thereof represent windows into the Hallandian orogenic evolution of Baltica. 2.2. Geology of the Romeleåsen horst The southernmost exposure of the Baltic Shield is delimited by the Tornquist zone; a major NW–SE trending steep fault zone

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Table 1 U–Pb zircon ages of Hallandian high-grade metamorphism across the Sveconorwegian and Blekinge-Bornholm provinces. ID in Fig. 2

Province/Tectonic unit/Locality

Blekinge-Bornholm province Unveined grey gneiss (Kullerön/84086) 1 2 Stromatic migmatitic gneiss (Lindö ÅJ01:21) 3 Migmatitic paragneiss (Nöteboda/C49) 4 Migmatitic paragneiss (Nöteboda/C50) 5 Migmatitic paragneiss (Nöteboda/C51) Sveconorwegian province Eastern marginal part of the Eastern Segment: low grade Sveconorwegian imprint Weakly gneissic, unveined porphyritic granite 6 (Apelskift/T0416) 7 Weakly gneissic, unveined porphyritic granite (Nissastigen/DC98-13) 8 Isotropic red leucocratic granite (Skinnarebo/SK0411K) 9 Weakly gneissic granite (Snyggebo/T0415) 10 Felsic band in migmatitic gneiss (Nissastigen/DC98-11) 11 Weakly gneissic granite (Västra Jära/T0401) 12 Cordierite-bearing migmatitic granofels (Nygård/JU12) Cordierite-bearing migmatitic granofels (Nygård/JU13A) 13 Cordierite-bearing migmatitic granofels (Nygård/JU13B) 14 Unmetamorphosed porphyritic granite 15 (Veberöd/MGO080025)

Na

Ea

Age (Ma)b

Reference

6223791 6219452 6167507 6167651 6167701

486440 521478 444996 444999 444908

c. 1440 c. 1440 1450–1420 1450–1420 1450–1420

Johansson et al., 2006 Johansson et al., 2006 Bogdanova et al., 2014 Bogdanova et al., 2014 Bogdanova et al., 2014

6412639

438690

1450–1350

Brander et al., 2012

6401749

434036

1366 ± 38

Brander et al., 2012

6397974 6410432 6401749 6402461 6160600 6160600 6160600 6167320

443716 435780 434036 438239 407290 407290 407290 402080

1378 ± 28 1426 ± 26c 1443 ± 9 1520–1380 1448 ± 9 1447 ± 5 1467 ± 7 c. 1450

Brander et al., 2012 Brander et al., 2012 Brander et al., 2012 Brander et al., 2012 This study This study This study This study

426112 418362 423902 419064 425544 363011 350064 349668 368535 371075 372092 355183 354184 381889 368291

1490–1360 1433 ± 6 1437 ± 6 c. 1440 1444 ± 15 1397 ± 4 1428 ± 9 1419 ± 12 1449 ± 11 1464 ± 8 1425 ± 7 1426 + 9/−4 1438 + 12/−8 1421 ± 7 1426 ± 18

Brander et al., 2012 Brander et al., 2012 Brander et al., 2012 Brander et al., 2012 Brander et al., 2012 Rimˇsa et al., 2007 Söderlund et al., 2002 Söderlund et al., 2002 Söderlund et al., 2002 Söderlund et al., 2002 Möller et al., 2007 Christoffel et al., 1999 Christoffel et al., 1999 Lundqvist et al., 2007 Austin Hegardt et al., 2005

368291

1415 ± 15

Austin Hegardt et al., 2005

High-grade parts of the Eastern Segment: intense penetrative high P/T Sveconorwegian reworking Weakly gneissic granite (Nybygget/T0408) 6412598 16 17 Unveined granitic gneiss (Marbohemmet/T0412) 6415654 18 Isoclinally folded migmatite leucosome (Vråna/DC98-14b) 6411274 19 Quartz-monzonitic gneiss (Flårekulla/T0413) 6418630 20 Gneissic granite (Kvistaberg/T0409) 6416116 21 Charnockitization of orthogneiss (Söndrum/CH, CZ, GN) 6280424 22 Unveined orthogneiss (Dagsås) 6329188 23 Veined orthogneiss (Dagsås) 6328833 26 Vein in veined orthogneiss (Gällared N/Vein sample) 6335053 27 Veined orthogneiss (Gällared S) 6331535 28 Veined orthogneiss (Högabjär/HB2) 6292797 29 Pegmatitic neosome (Särdal/Sample 3) 6293458 30 Mafic orthogneiss (Gåsanabbe/Sample 5) 6297624 31 Veined gneissic granite (Hedhuset/TEN050212) 6253009 6396889 32 Polymetamorphic migmatitic granodiorite hosting Sveconorwegian eclogite (Viared/EAH0207) 6396889 33 Polymetamorphic pegmatite granite hosting Sveconorwegian eclogite (Viared/DC03-116) a b c

c

UTM coordinate system (SWEref). Errors given as 2 or in the 95% confidence interval. One analysis only (207 Pb/206 Pb age).

(Fig. 1). The Romeleåsen horst (Figs. 1–3) was raised by fault reactivation in the Tornquist zone during Mesozoic and Tertiary compression (Bergerat et al., 2007). It exposes Precambrian bedrock, principally red to greyish red orthogneisses (Erlström et al., 2004). Modern age determinations are rare; U–Pb zircon analysis of granitic gneiss at Stenberget gave a poorly constrained protolith age of about 1.68 Ga (Johansson et al., 1993). This age falls in the 1.74–1.66 Ga age range typical of rocks of the Transscandinavian Igneous Belt (including the Eastern Segment of the Sveconorwegian province; e.g., Petersson et al., 2013, and references therein). The gneisses at Romeleåsen are granitic, finegrained and generally leucocratic, with variably developed gneissic foliation. The main minerals are generally quartz, perthitic microcline and plagioclase with only a few percent biotite and other dark minerals (ilmenite and Fe-oxide; Hjelmqvist, 1934). Aluminous grey gneiss with sillimanite, cordierite and/or garnet occurs in the central part of the Romeleåsen horst (Fig. 3). It has been interpreted as being of supracrustal origin, either volcanic or sedimentary (Hjelmqvist, 1934). Presumed metavolcanic rocks are granoblastic and dominated by quartz, microcline and plagioclase with ≤10% biotite, and small amounts of garnet, muscovite, and opaques. Presumed metasedimentary rocks are quartz phyllite, cordierite-sillimanite gneiss, garnet-muscovite gneiss and

garnet-biotite gneiss. Cordierite-sillimanite gneiss is often finegrained with blue-grey pseudomorphs after cordierite formed during retrogression. The cordierite-sillimanite gneiss is variably deformed and occurs as granofelsic to banded or foliated, and is in places migmatised. Several varieties of meta-mafic rocks occur. The gneisses in the northcentral part of the Romeleåsen horst contain concordant layers of amphibolite; they are probably co-eval with their host rocks. Other amphibolitic dykes cut the ductile structures and are younger (Erlström et al., 2004). Hjelmqvist (1934) identified several petrologically different types of amphibolite: cummingtonite amphibolite, hypersthene amphibolite, anthophyllite amphibolite and quartz amphibolite. Young “granites” occur locally, e.g., the reddish Romele granite (Veberöd, Stenberget, Nygård and north of Romeleklint, Fig. 3) and the greyish hornblende-bearing granodiorite at Beden (Fig. 3). The Beden granodiorite has been dated at 1464 ± 9 Ma (U–Pb zircon, Petersson et al., 2013), an age similar to the igneous rocks in the Blekinge-Bornholm province (Fig. 2; e.g., Zarin¸sˇ and Johansson, 2009 and references therein). In bedrock maps of the Romeleåsen horst, the Romele granite has been considered to be of the same age (Erlström et al., 2004). Similar to other presumably young granites at the Romeleåsen horst, it shows weak deformation (Hjelmqvist,

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Fig. 3. Geological map of Romeleåsen, revised after SGU bed rock map data base. Inset: Romeleklint (after Hjelmqvist, 1934).

1934). Other young intrusive rocks are greenish grey-reddish grey gneissic granite at the north-western tip of the Romeleåsen host (Billebjär), and medium-grained purplish-brown syenite in the central part, with 70–80% K-feldspar and plagioclase (cf. Fig. 3). Hypersthene-bearing NNE-trending metadolerite dykes (“hyperite dykes”) occur in the northcentral part of the Romeleåsen horst. They are fine- to medium-grained and black-brownish violet. The texture is ophitic to subophitic with plagioclase laths and augitic pyroxene grains partly altered to hornblende. U–Pb analysis of baddeleyite from similar metadolerites gave ages of c. 1.2 Ga (Cederberg, 2011). They may be associated with 1.2 Ga dolerite dykes along the Sveconorwegian Front farther north (Johansson and Johansson, 1990; Söderlund et al., 2005).

The Romeleåsen horst also exposes a dense NW-NNW trending steeply dipping unmetamorphosed dolerite dyke swarm emplaced during Carboniferous and Permian rifting (Erlström et al., 2004).

3. Petrography and mineral chemistry 3.1. Investigated localities This study presents data from four localities with exposures of aluminous grey gneisses at Nygård, Stenberget, Romeleklint and Veberöd (Fig. 3). Metadolerite and granitic gneiss are exposed at the same localities, with the exception of Nygård.

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3.1.1. Nygård (SWEREF: 6160600/407290) Nygård (30 × 30 m outcrop) consists of a homogenous fine- to medium-grained yellowish grey cordierite-bearing migmatitic granofels with bluish grey specks, a few mm–cm in size (Fig. 4a and b). The specks are pseudomorphs after cordierite. Faint leucosomes demonstrate partial melting (Fig. 4a). 3.1.2. Stenberget (SWEREF: 6158300/406400) Stenberget is an active quarry dominated by red mediumto fine-grained granitic gneiss, commonly with abundant coarsegrained red microcline. The eastern wall exposes several varieties of mostly well foliated fine-grained, locally garnetbearing and biotite-rich grey gneisses (presumably of supracrustal origin, Fig. 4d). Metadolerite sheets and dykes, folded, pinchand-swell shaped or boudinaged, are present in the granitic gneisses. 3.1.3. Romeleklint (SWEREF: 6165130/402500) Romeleklint was mapped in detail by Hjelmqvist (1934; Fig. 3). The largest outcrop consists of foliated metadolerite (amphibolite; Fig. 4f) cut by metadolerite (hyperite) dykes. Three smaller outcrops west thereof expose different varieties of aluminous gneiss and a larger eastern outcrop is made up of aluminous gneiss and metadolerite. Most gneisses are well foliated. Locally, leucosome with centimetre-sized garnet porphyroblasts is present. 3.1.4. Veberöd (SWEREF: 6167320/402080) An active quarry 2 km WNW of Veberöd exposes grey, finegrained, finely banded to laminated, felsic gneisses intercalated with reddish veins of coarse- to medium-grained K-feldspar-rich red granite (Romele granite, Fig. 4c and e). The granite is in places observed as several metres large bodies. Gneissic structures strike N and dip steeply W. A few metres wide black metadolerite dyke strikes N and dips c. 30◦ W. Structurally young, discrete, N-NNW-striking muscovite and chlorite-rich zones, presumably Sveconorwegian in age, are well-developed along the contact to the mafic dyke. 3.2. Investigated samples Eight samples of Crd-Grt-bearing granofels, 20 samples of Grtbearing Sil-Bt-gneiss and one sample of amphibolite have been examined. Mineral compositions were determined on carbon-coated thin sections using scanning electron microscopy (Hitachi S-3400N) fitted with an EDX analyser (Oxford Instruments with INCA software) at the Department of Geology, Lund University. Acceleration voltage was 15 kV, the working distance 10 mm, and spot size <1 ␮m. Analysis was performed with a live-time of 60–80 s and calculated using natural and synthetic standards; analytical uncertainty is <2%. Line-scans were performed for characterisation of compositional zoning within single grains. Representative compositions for minerals in Crd-Grt-bearing granofels (samples JU12 and SB6B), Grt-bearing Sil-Bt gneiss (samples JU25A2, JU8B and JU25B) and amphibolite (sample JU9B) are presented in Tables 2–4. 3.2.1. Aluminous gneisses The aluminous gneisses are cordierite-garnet-bearing migmatitic granofels and garnet-bearing biotite-sillimanite gneiss, the latter generally foliated. 3.2.1.1. Cordierite-garnet-bearing granofels. The matrix of the cordierite-garnet granofelses is granoblastic with mm-sized grains of quartz, K-feldspar and plagioclase. Centimetre-sized yellowish or

7

greenish spongy aggregates in the matrix (Fig. 5a and b) are pseudomorphs after cordierite, and occasionally contain remnants thereof. Numerous mm-sized opaque grains and a few biotite grains are also present. Quartz is weakly undulose and K-feldspar is crypto- to microperthitic (Or85–95 ). Plagioclase is An25–35 . Cordierite has been a major phase but is only preserved as rare remnants in pseudomorphs (Fig. 5c); Fe/(Fe + Mg) = 0.39–0.44. Garnet grains of different sizes (mm–cm) are in places spatially associated with cordierite pseudomorphs (Fig. 5d). Smaller garnet grains are sub- to euhedral and in many places contain inclusions of biotite and quartz. Locally, garnet has inclusions of spinel, staurolite and/or muscovite. Some inclusions are composite aggregates of biotite, Fe-oxide, spinel and fine-grained intergrowths of Fe-oxide + corundum and ilmenite + Ti-oxide + quartz (Fig. 5e). Resorption of garnet is shown by tails of sillimanite and fine-grained aggregates of sillimanite, spinel and Fe–Tioxides. Garnet composition is generally homogeneous (Fig. 5f), almandine-rich (XAlm = 0.70–0.75) with low amounts of pyrope (XPrp = 0.1–0.2) and spessartine (XSps = 0.08–0.10) and minor grossular (XGrs = 0.02–0.03). One sample from Stenberget is richer in spessartine at the expense of almandine. Locally along rims and cracks there is an increase in almandine, grossular and spessartine and decrease in pyrope. Biotite occurs as <1 mm long flakes in the matrix, in many places in aggregates with sillimanite, Fe-oxide and locally minor spinel, and as inclusions in garnet. In pseudomorphs after cordierite, biotite forms a corona around Fe-oxides. Fe/(Fe + Mg) = 0.5–0.6 in the matrix, decreasing to 0.40–0.45 close to garnet and in inclusions in garnet. Ti is 0.15–0.25 cations per formula unit (based on 11 O). Symplectites of biotite and plagioclase are also observed. Greenish biotite (Fe/(Fe + Mg) = 0.5 and Ti absent or very low) occurs in the pseudomorphs after cordierite; it has also been observed as a corona around garnet. Prismatic sillimanite occurs in aggregates with biotite. In cordierite pseudomorphs sillimanite is needle shaped or fibrolitic. It is nearly pure Al2 SiO5 with 0.4–1.0 wt.% Fe. Spinel is found locally as single-phase inclusions in garnet, or in composite inclusions together with biotite and intergrowths of hematite + corundum + FeTi-oxide aggregate, the latter consisting fine-grained ilmenite, Ti-oxide, and quartz (Fig. 5e). Spinel can also be found in the matrix associated with biotite, sillimanite, Fe-oxide and ilmenite with fine-grained Ti-oxide and quartz. Spinel is hercynite with variable amounts of gahnite (ZnAl2 O4 , <25 mol%) and small amounts of galaxite (MnAl2 O4 ). Staurolite has been found in one sample from Stenberget as an inclusion in garnet (Fig. 6c); it contains about 5 mol% of the Zn endmember. Small amounts of Mn are present. Muscovite occurs in small flakes together with green biotite in pseudomorphs after cordierite. Accessory minerals include abundant rounded zircon grains 20–100 ␮m long. Fe-oxides are hematite (calculated as 98.5% Fe2 O3 based on total weight-% Fe), often associated with biotite and spinel. Ilmenite contains <30 mol% pyrophanite (MnTiO3 ). Monazite and pyrite are present. 3.2.1.2. Garnet-bearing biotite-sillimanite gneiss. Garnet-bearing biotite-sillimanite gneisses (Fig. 6a and b) have a foliation and/or lineation defined by biotite and sillimanite anastomosing around porphyroclastic porphyroblasts of partly resorbed garnet. In some samples greenish aggregates are present, interpreted as pseudomorphs after cordierite. Quartz forms irregular elongate undulose grains 5–10 mm long and also fine recrystallised grains. Plagioclase is An25–35 , commonly clear with some seritization. K-feldspar is Or90–95 and cryptoperthitic.

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Fig. 4. Rocks from Romeleåsen. (a) Cordierite-bearing migmatitic granofels, Nygård. Leucosome dashed red. Compass, 6 cm, for scale. (b) Cordierite-bearing granofels, Nygård (JU12). (c) Finely banded felsic gneiss with K-feldspar-rich granite veins, Veberöd quarry. Coin, 20 mm, for scale. (d) Garnet-bearing biotite-sillimanite gneiss, Stenberget (SB1AJU). (e) K-feldspar porphyritic granite, Veberöd quarry (MGO080025). Coin 24 mm for scale. (f) Amphibolite, Romeleklint (JU9B). (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

Biotite forms irregular flakes 0.1–1.5 mm or fine recrystallised grains. Fe/(Fe + Mg) = 0.5–0.6, Ti = 0.2–0.5 cations per formula unit (based on 11 O). Green biotite is present in pseudomorphs after cordierite; in these grains Ti is absent or very low. Garnet grains vary in size from 0.5 to a few millimetres. Large grains are often fractured and contain inclusions of quartz, biotite,

sillimanite needles and opaques. Small garnet grains are euhedral to subhedral and also have inclusions of quartz, biotite, sillimanite and opaques. Compositional profiles are flat with Alm70–80 Prp7–15 Grs2–8 Sps5–20 (Fig. 7a and b). Sillimanite occurs in fibrolite bundles and as prismatic grains 50–200 ␮m long, and also as inclusions in garnet (Fig. 6d). It contains 0.4–1.0 wt.% Fe.

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Table 2 Representative EDS microprobe data for minerals in cordierite-garnet bearing granofelses. Mineral Sample Analytical site

Grt JU12 9035 (n = 14)

Bt JU12 9035 Grt incl (n = 9)

Bt JU12 9035 Matrix (n = 5)

Pl JU12 9035 Matrix (n = 9)

Crd JU14 8020 Matrix (n = 2)

SiO2 TiO2 Cr2 O3 Al2 O3 FeO tot MnO MgO CaO Na2 O K2 O V2 O5 ZnO Total

36.14

35.49 3.33

34.58 3.47

58.34 0.05

46.64

18.75 16.82

17.73 20.99 0.09 9.22

25.31

32.40 8.85 0.57 7.43

O for calca Si Ti Al Cr Fe2+ Fe3+ Mn Mg Ca Na K Zn Total cations Fe/(Fe + Mg) XFe XMg XCa XMn XZn An Ab a b

21.13 32.25 4.14 3.52 0.95

12.24

9.71

98.12

96.22

95.78

98.58

95.89

12 2.96

11 2.65 0.19 1.65

11 2.65 0.20 1.61

8 2.63 0.00 1.35

18 4.94

1.05 0.00

1.35 0.00 0.01 1.06

2.16 0.05 0.29 0.43 0.08

1.36

0.91

0.36 58.19 33.04 4.28

Spl JU12 9535 Matrix (n = 2)

58.17 35.51 1.19 2.79

Ilm JU12 9535 Matrix (n = 2)

Rt JU12 9535 Matrix (n = 2)

Hem JU12 9535 Matrix (n = 3)

St SB6B 8030 Grt incl (n = 5)

97.08

0.12 0.15

27.58

51.94

33.47 0.00

0.17 34.20 13.27

3.61

0.50 97.59b

56.31 12.68 0.82 2.06

19.71 19.99

6.72 7.99 0.08

9.52 0.06

2.04

Spl JU12 9035 Grt incl (n = 3)

Bt SB6B 4050 pseudo (n = 1)

10.99

8.44 0.22

0.95

8.00

7.80

7.82

0.84 0.73 0.14 0.03 0.10

0.44

0.56

4.05 0.72 0.07 0.05 1.17

2.88 98.76

4.11 101.8

99.58

100.69

4

4

3

2

1.94

0.99 0.01

0.98

1.96 0.01 0.79 0.01

0.72 0.01 0.28

0.04

98.55 3

0.01

1.41 100.87 23.5 3.81 9.17

0.18

0.84 0.01 0.03 0.12

1.46

0.06 3.01

0.09 3.02

0.81 0.75 0.18

0.88 0.77 0.11

0.78 0.69 0.20

0.07

0.03 0.08

0.04 0.07

1.97 0.09 0.42

0.33 0.70 0.00 5.03

92.60 11 2.60 0.00 1.81 1.10 0.20 1.27

0.84 11.00 0.40

2.01

1.02

2.00

0.14 15.10

7.82 0.46

0.32 0.68

Cation compositions include a stoichiometric estimate of Fe3+ calculated with the AX software (Holland, 2012). Fe2 O3 .

Spinel is found locally in two different textural associations: (1) as inclusions in garnet, and (2) in the biotite-rich matrix, surrounded by a corona of sillimanite. In the latter association, spinel is intergrown with very fine-grained staurolite and corundum (Fig. 7c and d). Here spinel is very Zn-rich (≤60 mol%). Fine-grained staurolite in association with spinel is Zn-rich (≤17 mol%). Muscovite occurs mainly as fine grains in aggregates, occasionally as larger flakes ≤0.5 mm. Accessory minerals include ilmenite (with ≤20 mol% MnTiO3 ), Ti-oxide (together with corundum as a late replacement after ilmenite), hematite, zircon and monazite. 3.2.2. Amphibolite Amphibolite (sample JU9B) is dark, fine-grained, foliated, and contains cummingtonite. The matrix is equigranular but a spaced foliation is defined by finer-grained aggregates of amphibole and biotite. Amphiboles are hornblende and cummingtonite. Hornblende forms olive-green to brownish anhedral grains, in places twinned. It is magnesiohornblende to edenite with Fe/(Fe + Mg) = 0.42–0.49. Cummingtonite forms colourless anhedral aggregates as pseudomorphs after orthopyroxene with Fe/(Fe + Mg) = 0.43–0.45. Biotite forms reddish brown to light brown flakes. Fe/(Fe + Mg) is 0.39–0.45.

Garnet grains are few, about 0.5 mm large, sub- to anhedral, and have inclusions of amphibole, plagioclase, biotite and ilmenite. The composition is Alm65–70 Prp10–20 Grs10–15 Sps5 . Cores are homogeneous in composition, whereas rims are locally higher in almandine and spessartine and lower in pyrope and grossular (Fig. 7e and f). Plagioclase is clear, in places with slightly curved twins. The composition corresponds to labradorite (An50–70 ), however close to apatite it is andesine (An30–50 ). Isolated remnants of orthopyroxene have been observed. Accessory minerals include ilmenite, apatite and dolomite. Dolomite is most likely the result of crystallisation from secondary fluids. 4. Petrological interpretation 4.1. Parageneses and reactions The aluminous gneisses (cordierite-garnet-bearing granofelses and garnet-bearing biotite-sillimanite gneisses) from Romeleåsen have similar bulk compositions (Table 5) and are therefore expected to have similar modal contents of main minerals if equilibrated at the same metamorphic grade. As is seen from the petrographic descriptions above, modal contents of minerals vary. Textural relations, in particular the resorption of porphyroclastic garnet in

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Table 3 Representative EDS microprobe data for minerals in the biotite-sillimanite-garnet gneiss. Mineral Sample Analytical site

Grt JU25A2 2535 Grt core (n = 9)

Bt JU25A2 2535 matrix (n = 6)

Pl JU25A2 2535 matrix (n = 5)

SiO2 TiO2 Cr2 O3 Al2 O3 FeO tot MnO MgO CaO Na2 O K2 O ZnO Total

36.40

34.21 3.18

57.65

20.96 34.26 2.95 3.37 1.02

17.96 18.50

25.03

98.95

92.98

97.49

O for calca Si Ti Al Cr Fe2+ Fe3+ Mn Mg Ca Na K Zn Total cations

12 2.96

11 2.67 0.19 1.65

8 2.64

4

1.35

1.95 0.02 0.36 0.02

Fe/(Fe + Mg) XFe XMg XCa XMn XZn An Ab a

9.41

Spl JU25A2 2575 (n = 2)

St JU25A2 2575 (n = 3)

Bt JU8B 0540 pseud (n = 6)

St JU8B 0540 pseud (n = 2)

St JU25B 0590 Grt incl (n = 4)

26.02 0.44

37.13

27.79

27.65

55.51 14.63 0.72 2.42

53.60 10.93

21.19 18.40

55.03 12.61

1.53

10.84

55.01 11.77 0.65 1.08

25.31 98.23

3.13 95.65

97.62

2.66 98.97

1.08 98.67

23.5 3.79 0.05 9.21

11 2.73

23.5 3.91

23.5 3.88

1.84

9.14

9.11

1.33

1.13 0.00

1.39

1.48

0.11

0.33

1.19

0.23

0.45

0.57 3.03

0.34 15.05

0.30 15.02

0.11 15.05

0.76 0.35 0.10

0.67 0.17

0.86 0.70 0.12

0.77 0.72 0.22

0.55

0.17

0.04 0.14

0.05

9.72

2.01 2.26 0.07 0.20 0.41 0.09

2.17

7.11 7.69 10.06

1.21 0.00 1.10 0.35 0.68 0.97

8.00

7.79

0.85 0.76 0.14 0.03 0.07

0.52

0.94 5.02

7.83 0.49

0.34 0.66

Cation compositions include a stoichiometric estimate of Fe3+ calculated with the AX software (Holland, 2012).

biotite-sillimanite gneisses, suggest significant retrograde reequilibration in association with deformation at biotite-sillimanite grade. It is also possible to find textures and assemblages transitional between typical granofels and biotite-sillimanite gneiss. Sample JU8B (Fig. 6a) contains greenish areas identified as pseudomorphs after cordierite, while sample JU4 (Fig. 6b) has strongly resorbed and fractured garnet grains and biotite-fibrolite bundles are present. The granofelses are markedly richer in K-feldspar and pseudomorphs of cordierite, and their modal contents of sillimanite and biotite are low (c. 3 modal-% biotite, the green variety in cordierite pseudomorphs excluded). This suggests that the granofels assemblage represents P–T conditions close to the metamorphic peak, probably not far below biotite-out temperatures. Remnants of early minerals include staurolite, sillimanite and spinel. Spinel occurs in small amounts in the aluminous gneisses, in a few different textural associations. It can be found as single grain inclusions in garnet in granofels and gneiss (Fig. 5d). It can also be found in the matrix of granofels associated with biotite, sillimanite, Fe-oxide and ilmenite with fine-grained Tioxide and quartz. Finally it is found in the gneisses intergrown with very fine-grained staurolite and corundum in biotite-rich matrix, surrounded by a corona of sillimanite (Fig. 7c and d). Spinel is a hightemperature mineral in metamorphic rocks (Bowles et al., 2011), but Zn increases its stability field towards lower temperatures and higher pressures (Hand et al., 1994). The Zn-content of spinel in the Romeleåsen aluminous gneisses is high, corresponding to a 6–60% gahnite. Zn may have been provided from the breakdown of minerals such as biotite, staurolite, or through desulphurisation

of sphalerite (ZnS; Bowles et al., 2011). Staurolite with observed high Zn-content may be a candidate for formation of spinel. The association of spinel with staurolite and corundum in sillimanite coronas may have not been part of the equilibrium assemblage, but may have developed locally in Al-rich and Si-undersaturated subdomains. Sillimanite inclusions in garnet are common and could have formed before or concomitantly with garnet growth, e.g., via the discontinuous reaction (McLellan, 1985): St + Ms + Qtz = Bt + Grt + H2 O

(1)

This reaction is the classical staurolite-out reaction in metapelitic rocks and can also explain staurolite inclusions in garnet. Partial melting has occurred in both cordierite-bearing granofels and garnet-bearing sillimanite-biotite gneiss. This has resulted in the local formation of poikilitic garnet porphyroblasts in leucosome, suggesting water-undersaturated melting by breakdown of biotite according to the reaction: Bt + Pl + Qtz + Sil = Grt + Kfs + L

(2)

This reaction takes place at temperatures above 650–700 ◦ C (Vernon and Clarke, 2008). The major cordierite producing reaction is likely: Sil + Bt + Qtz = Crd + Kfs + L

(3)

Fine-grained symplectite of biotite and plagioclase are also indicative of presence of melt (Sawyer, 2008). Observed foliations

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Table 4 Representative EDS microprobe data for minerals in amphibolite. Mineral Sample Analytical site

Grt JU9B 8025 Grt core (n = 8)

Grt JU9B 8025 Grt rim (n = 2)

Bt JU9B 8025 Matrix (n = 6)

Pl JU9B 8025 Matrix (n = 4)

Mg-Hbl JU9B 8025 Matrix (n = 4)

Cum JU9B 8025 Matrix (n = 4)

En JU9B 1070 (n = 4)

SiO2 TiO2 Al2 O3 FeO tot MnO MgO CaO Na2 O K2 O Total

36.84

36.35

57.91

50.08

20.87 31.48 2.43 3.05 4.28

44.34 1.05 10.27 15.31

51.24

20.93 29.83 2.02 3.57 5.47

35.78 3.73 15.69 17.30

1.11 24.02 0.60 17.01 0.63

0.81 30.00 0.73 17.09 0.40

98.66

98.45

94.33

99.12

12 2.97

12 2.93

23 7.61

6 1.91

1.99 1.93 0.08 0.14 0.43 0.47

1.99 2.13 0.08 0.17 0.37 0.37

0.04 0.96 0.10 0.02 0.97 0.02

8.00

8.03

2.59 1.77 0.40 0.09 15.49

0.20 2.98 0.38 0.08 3.76 0.10

15.11

4.02

0.82 0.65 0.14 0.16 0.05

0.85 0.70 0.12 0.12 0.05

0.37

0.44

0.50

a

O for calc Si Ti Al Fe2+ Fe3+ Mn Mg Ca Na K Total

Fe/(Fe + Mg) Alm Prp Grs Sps An Ab a

25.37

12.52

11.55 10.97 1.36 0.45 95.31

6.85 8.10 9.49 94.52

98.23

11 2.70 0.21 1.40 1.09 0.08

8 2.63

23 6.66 0.12 1.82 1.54 0.39

1.36

1.41 0.33 0.71 0.92 7.81

5.04

0.44

0.32 0.68

Cation compositions include a stoichiometric estimate of Fe3+ calculated with the AX software (Holland, 2012).

and lineations defined by biotite and sillimanite are the result of the breakdown of garnet and/or cordierite, i.e., the reverse of Reactions (2) and (3) above. The flat compositional profiles of garnet in the aluminous gneisses reflect homogenisation by intracrystalline diffusion at high temperatures (>650 ◦ C) during prograde and peak regional metamorphism (cp. Kohn, 2003). The small-scale zoning at rims, with increasing almandine and corresponding decrease in pyrope, is due to either retrograde net transfer reactions, diffusional Fe–Mg exchange, or both. Increasing Mn and Ca contents are indicative of net transfer reaction as these elements are partitioned back into the reacting garnet (unless Ca- and/or Mn-bearing products are formed). The amphibolite contains a few relicts of orthopyroxene (enstatite-hypersthene) and it is difficult to decipher whether they

are remnants of primary igneous phases or metamorphic. The rare occurrence of better-preserved, clear orthopyroxene grains with a partial corona of cummingtonite suggests retrograde breakdown of orthopyroxene. The overall flat compositional profiles of garnet in amphibolites suggest homogenisation at high metamorphic temperatures. Decrease in Ca content along the garnet rims may be due to resorption and formation of plagioclase and/or amphibole during retrogression. Pseudomorphs after cordierite are ascribed to low-temperature retrograde breakdown of cordierite, commonly termed pinitization. Ogiermann (2003) described different pinitization processes, and distinguished four pinite types. B-type (border) consists of muscovite and green biotite (low titanium) that form through a reaction of cordierite, K-feldspar and water. M-type (mat) consists of very fine-grained aggregates of muscovite and chlorite and

Table 5 Bulk rock ICP analyses (wt.%) of samples from Romeleåsen. Grt-Bt gneisses

SiO2 TiO2 Al2 O3 Cr2 O3 Fe2 O3 MnO MgO CaO Na2 O K2 O P2 O5 LOI Sum

Crd-Grt bearing granofels

Amphibolite

JU4

JU8B

JU25Ab

JU25B

SB1AJU

JU12

SB6

JU9B

65.85 0.58 17.93 0.003 5.03 0.13 1.80 1.34 2.30 2.93 0.04 1.8 99.73

68.02 0.53 16.25 0.003 4.34 0.12 1.23 1.94 2.44 3.88 0.07 0.8 99.62

66.56 0.66 17.51 0.004 5.29 0.14 1.46 1.28 1.61 3.58 0.07 1.5 99.66

72.47 0.46 15.14 <0.002 3.19 0.07 0.88 2.15 2.33 2.00 0.05 1.0 99.74

58.64 0.93 23.21 0.004 7.69 0.23 1.79 0.59 0.81 1.69 0.03 4.2 99.81

69.92 0.54 15.47 <0.002 4.35 0.11 0.83 2.08 2.34 2.23 0.05 1.8 99.72

64.69 0.76 18.30 0.004 4.93 0.14 1.36 1.31 3.19 3.42 0.04 1.6 99.74

46.25 1.68 15.96 0.008 14.28 0.19 8.24 8.09 2.10 1.39 0.38 1.1 99.67

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Fig. 5. Mineral assemblages and microstructures of cordierite-garnet-bearing granofels. (a) Cordierite-garnet-bearing granofels from Nygård (JU12). (b) Cordierite-garnetbearing granofels from Stenberget (SB6). (c) Photomicrograph (PPL) of pseudomorph after cordierite. (d) Photomicrograph (PPL) showing garnet in (a) with spinel inclusions. Position for zoning profile marked in red. (e) Back-scattered electron (BSE) image of a composite inclusion in garnet in (d) of spinel, biotite, hematite, corundum and ilmenite with the latter mineral partly replaced by Ti-oxide and quartz. (d) Compositional zoning of garnet in (d) from right to left. Fe/(Fe + Mg) ratios and mole fractions of almandine (Alm) on left axis, pyrope (Prp), grossular (Grs) and spessartine (Sps) on right axis. Point 4 and 5 at spinel inclusion. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

can also contain clays. F-type (fissure) and i-type (isotrope) pinite are enriched in Ca and can contain amorphous material. All four types can be observed in the cordierite pseudomorphs from Nygård (Fig. 5c). Fluid is required for the supply of potassium and water for the reactions.

Ti-oxide and quartz are abundant in association with partly resorbed and altered ilmenite (Fig. 5e). Rutile is usually indicative of medium- to high-grade metamorphism, and in particular of high pressures (Meinhold, 2010); however, other Ti-oxides such as anatase and brookite can form under low- and medium-grade

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Fig. 6. Mineral assemblages and microstructures of garnet-bearing biotite-sillimanite gneiss and Romele granite. (a) Garnet-bearing biotite-sillimanite gneiss from Romeleklint (JU8B). (b) Garnet-bearing biotite-sillimanite gneiss from Romeleklint (JU4). (c) BSE image of staurolite inclusions in garnet porphyroblast (SB6B). (d) BSE image of sillimanite inclusions in garnet porphyroblast (JU25B). (e) Photomicrograph (PPL) of Romele granite. Altered feldspar (top right) is saussuritized plagioclase (MGO080025). (f) Same as (e) with PPX showing tartar twinned microcline megacrysts surrounded, and partly intervened, by fine-grained domains of feldspar, quartz and myrmekite (MGO080025).

conditions. Diagenetic and low-grade replacement of ilmenite by Ti-oxide + silicates has been documented e.g., in metasedimentary rocks (Luvizotto et al., 2009). Ti-oxide can also form under lowgrade conditions through oxidation and metasomatism of ilmenite (Putnis, 2009; Engvik et al., 2011).

4.2. Peak paragenesis The peak paragenesis for the cordierite-garnet bearing granofels is garnet + sillimanite + cordierite ± Ti-rich biotite + melt + plagioclase + K-feldspar + quartz + ilmenite. Since intermediates

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between the granofels and the garnet-bearing biotitesillimanite gneiss containing pseudomorphs after cordierite can be observed, it is assumed that the gneiss had the same peak paragenesis. The peak paragenesis for the amphibolite is hornblende + plagioclase + biotite + garnet + ilmenite ± quartz ± orthopyroxene.

5. Phase equilibria modelling Pseudosections (or isochemical phase diagrams) show all the stable multivariant phase assemblages in a chosen chemical system for a specified rock composition and can be used to constrain P–T conditions and elucidate metamorphic changes (Powell and Holland, 2010 and references therein). Use of P–T pseudosections can be a better alternative than geothermobarometry for determination of the P–T conditions of metamorphism, as P–T pseudosections not are dependent on preserved original mineral compositions (Powell and Holland, 2008). Conventional geothermobarometry, using e.g., the garnet-biotite thermometer and the GBPQ/GASP barometer (Berman, 1988, 1991, 2007) on the aluminous gneisses presented here, resulted in temperatures and pressures (550–650 ◦ C and 2–4 kbar) that are far too low for the observed peak paragenesis. Equilibrium phase diagrams using the analysed bulk rock compositions were calculated in the system MnNCKFMASHTO using THERMOCALC v3.33 (Powell and Holland, 1988 update October 2009) with the thermodynamic database of Holland and Powell (1998) updated to dataset 5.5, created 22 November 2003 (Powell and Holland, 2012). The phases considered in the modelling and solution models used are garnet, biotite, ilmenite and hematite (White et al., 2005), plagioclase-K-feldspar (Holland and Powell, 2003), orthopyroxene, magnetite and spinel (White et al., 2002), chlorite, cordierite and staurolite (Mahar et al., 1997; Holland and Powell, 1998), muscovite-paragonite (Coggon and Holland, 2002), silicate melt (White et al., 2007), epidote (Holland and Powell, 1998). Other phases and H2 O are considered as pure end-member phases. Bulk rock compositions were determined at ACME laboratories by inductively coupled plasma emission spectroscopy (ICP) on rock pulp from the same samples used for preparing thin sections and are presented in Table 5. As P is not included in the modelling, the CaO of apatite was subtracted from the total CaO of the rock. Cr2 O3 was disregarded when calculating the composition for modelling. The bulk composition formula Fe2 O3 was recalculated as Fe2+ . Ferric iron is present in almost all metamorphic rocks and may affect the phase relations (Diener and Powell, 2010), but it is commonly difficult to quantify and therefore not often included in modelling. In the rocks at hand hematite is present, which shows that the rocks are oxidised. Hjelmqvist (1934) analysed both Fe2+ and Fe3+ in aluminous gneisses from Romeleåsen and translated to oxygen ratio (Fe3+ /(Fe3+ + Fe2+ )); the result corresponds to a value of about 0.3. This should be a maximum value, because weathering processes may have oxidised the sample (cp. the low grade replacements of ilmenite, above), and therefore a value of 0.2 was chosen in the modelling. The aluminous gneisses show evidence of partial melting. The bulk composition therefore represents the final rock composition after possible loss of the melt and a pseudosection based solely on the analysed bulk composition may not be valid for the prograde evolution of the rock (e.g., White et al., 2004; Diener et al., 2008; Indares et al., 2008). The H2 O content of the sample was estimated so that the inferred peak metamorphic assemblage is stable directly above the solidus where the assemblage would have been in equilibrium with the last remnants of the melt (Diener et al., 2013). Pseudosections for compositions representing the final rock after

possible loss of melt are presented for the cordierite-garnet bearing granofels sample JU12 and the garnet-bearing biotite-sillimanite gneiss sample JU4 in Fig. 8a and b, respectively, in the P–T range 0.5–10 kbar and 500–900 ◦ C. The compositions used for calculating the pseudosections are given in Table 6. The topology of the pseudosections for samples of granofels and gneiss is similar, and dominated by 5- and 6-variance fields. In this P–T range quartz, plagioclase and ilmenite is everywhere present, K-feldspar everywhere except at very high T at low P, and magnetite is present below about 6 kbar (although hematite is observed in the samples). Garnet is stable above 3–4 kbar at temperatures above 650–700 ◦ C and also at lower pressures below this temperature. Cordierite is stable in the lower right part of the diagram, muscovite in the upper left, sillimanite in the mid and upper right, and biotite is stable from 500 ◦ C at low pressures to 800 ◦ C at low-intermediate pressure. Spinel becomes stable at temperatures above 800 ◦ C at mid-pressures. The cordierite-bearing granofelses from Nygård show lowstrain gneissic structures (Fig. 4a), whereas the garnet-bearing biotite-sillimanite gneisses from Romeleklint and Stenberget are heterogeneously deformed with locally strong foliations defined by biotite and sillimanite (Fig. 4d). As shown by Fig. 8a and b the topology of the pseudosections for the aluminous gneisses are similar, so if they have been exposed to the same conditions, they should contain the same assemblages. It is most likely that all parageneses had the assemblage cordierite-garnet as for Nygård at peak conditions, and that the (garnet)-sillimanite-biotite assemblages reflect re-equilibration during high- to medium-temperature retrogression. The peak assemblage garnet + sillimanite + cordierite ± Ti-rich biotite + plagioclase + K-feldspar + quartz + ilmenite + magnetite (hematite) + melt is represented by the narrow 3-variant field below the cordierite-in line and the biotite-out line where the solidus is joining. The metamorphic peak conditions for the cordierite-bearing granofels are about 700 ◦ C and 4 kbar while for the garnet-bearing biotite-sillimanite gneiss they indicate about 750 ◦ C and 5 kbar. The biotite-out line is, however, dependent on the modelling of Ti in biotite (Tajˇcmanová et al., 2009). The observed content of Ti (0.20 p.f.u. based on 11 O atoms) is significantly higher than predicted (0.04 p.f.u.) at the highest temperature limit of the biotite stability field. It is therefore probable that the stability field of biotite should be expanded to higher temperatures. The shape and position of the 3-variant field differ little between the granofels and gneiss. The choice of water content therefore results in an uncertainty as to where the solidus joins this field and in consequence the definition of peak conditions. As discussed earlier, the flat compositional profiles of garnet in the aluminous gneisses with minor retrograde zoning along rims are interpreted to reflect peak conditions. Although the isopleths of almandine and spessartine are compressed together in the 3-variant field, the observed 73% almandine and 9% spessartine intersect at P–T conditions of ∼700 ◦ C and ∼4.2 kbar. With the uncertainties described above, 700–750 ◦ C and 4–5 kbar is considered to represent peak P–T conditions. The pseudosection of the residiuum should theoretically be relevant to investigate retrograde processes. However, in the absence of fluids or melt the retrograde reactions will be sluggish and the metamorphic peak assemblage will be preserved (Guiraud et al., 2001; White and Powell, 2002). If fluids are present during retrogression, rehydration will occur, leading to new mineral assemblages (including e.g., micas). Retrograde reactions observed in the granofels and gneiss include medium- to high-temperature resorption of garnet associated with deformation and formation of fine-grained sillimanite and biotite, and low-temperature breakdown of cordierite through pinitization reactions to muscovite, biotite and chlorite.

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Fig. 7. Mineral assemblages and microstructures of garnet-bearing biotite-sillimanite gneiss and amphibolite. (a) BSE image of garnet (sample JU25Ab) with location of zoning profile in red. (b) Compositional profile of garnet in (a) from rim at quartz to rim at muscovite. Fe/(Fe + Mg) ratios and mole fractions of almandine (Alm) on left axis, pyrope (Prp), grossular (Grs) and spessartine (Sps) on right axis. (c) Photomicrograph (PPL) of composite spinel-rich aggregate surrounded by a corona of sillimanite in biotite-rich matrix (sample JU25Ab). (d) BSE close-up of the same domain as in (c), showing staurolite and corundum intergrowths with spinel, and the presence of late muscovite in biotite domains. (e) BSE image of garnet in amphibolite sample JU9B with location of zoning profile in red. (f) Compositional profile of garnet in (e) from rim at plagioclase to rim at cummingtonite. Fe/(Fe + Mg) ratios and mole fractions of almandine (Alm) on left axis, pyrope (Prp), grossular (Grs) and spessartine (Sps) on right axis. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

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Fig. 8. Calculated P–T pseudosections for the residuum composition of (a) cordierite-garnet-bearing granofels sample JU12 with isopleths (mol%) for almandine (red) and spessartine (blue), and (b) garnet-bearing biotite-sillimanite gneiss sample JU4 calculated with the bulk composition given in Table 6. Assemblages for the largest fields are indicated; all fields include quartz, plagioclase and ilmenite. The peak assemblage is indicated with a hatched circle. Arrows represent suggested P–T path. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

In an open system with melt loss it is generally not possible to recover a bulk composition that is representative of the exact protolith. However, forward modelling has shown rather small changes on the major sub- and supra-solidus topologies (White and Powell, 2002), and it is therefore possible to estimate a bulk composition that can be used for prograde modelling. In order to investigate the prograde evolution of the aluminous gneisses, model protolith compositions were estimated for samples JU12 and JU4 using melt-reintegration according to the method outlined in White et al. (2004) and Diener et al. (2013). This involved adding melt of equilibrium composition incrementally along an assumed prograde path starting at peak metamorphic conditions to lower P–T conditions until melt coexists with H2 O at the wet solidus. For samples JU12 and JU4 the melt reintegration followed backward along an inferred clockwise path peaking at 5 and 6 kbar, respectively, for the two samples. Melt added amounted to 11 mol% and 12 mol% for sample JU12 and JU4, respectively. The estimated protolith compositions for the two samples are presented in Table 6 and their calculated pseudosections in Fig. 9. The wet solidus for the two samples occurs at 650–670 ◦ C. Contours for modal proportions of melt are presented for sample JU12 in Fig. 9a, which shows that the melt production across the wet solidus is minor (1–3 mol%), and it is not until the 4-variant field Grt-Bt-Ms-KfsSil-Pl-Ilm-Qtz-Liq is reached that melt generation increases to 10 mol% over a very small temperature interval. The stability field for garnet starts at higher pressures in the subsolidus region of the

protolith diagrams compared to the residuum diagrams (Fig. 8). In Fig. 9 water is taken as in excess for the subsolidus part of the diagrams to simulate water liberation during dehydrating prograde reactions. Staurolite is present in the P–T range 3–5 kbar and 550–600 ◦ C for sample JU12 and in a larger field at higher pressure and temperature for sample JU4. Chlorite is present below c. 550 ◦ C and c. 6 kbar. In the subsolidus part muscovite is present almost everywhere and paragonite and epidote are also present at higher pressures. Spinel in these pseudosections is only found at temperatures at or above 800 ◦ C. As mentioned previously one explanation for the presence of spinel in the samples is that spinel was not part of the equilibrium assemblage but developed in Al-rich subdomains, another that the high Zn-content extends the stability field of spinel to lower temperature. 6. U–Th–Pb zircon geochronology Four samples were selected for U–Th–Pb SIMS zircon analysis. Two varieties of cordierite-garnet-bearing migmatitic granofels (JU12 and JU13) from Nygård (Fig. 3) were investigated with the aim to directly constrain the age of the high-temperature metamorphism, and to compare the zircon characteristics in the different samples. Sample JU12 is a patch migmatite with weakly segregated leucosome that could not be separated from the mesosome (Fig. 4a and b). This sample has the best preserved peak-temperature

Table 6 Bulk compositions (mol%) used to construct the pseudosections.

Fig. 7a (JU12) Fig. 7b (JU4) Fig. 8a and b (JU12) Fig. 8c (JU4) a

SiO2

TiO2

Al2 O3

FeO

MnO

MgO

CaO

Na2 O

K2 O

O

H2O

76.37 72.23 74.93 71.13

0.44 0.48 0.40 0.43

9.96 11.59 9.63 11.10

3.58 4.15 3.23 3.72

0.10 0.12 0.092 0.11

1.35 2.94 1.22 2.63

2.36 1.51 2.15 1.38

2.48 2.45 2.52 2.53

1.55 2.05 1.67 2.13

0.36 0.42 0.32 0.37

1.46 2.07 3.82a 4.48a

H2 O was taken as in excess for the subsolidus part of the pseudosection.

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mineral parageneses of the investigated samples and was used for pseudosection modelling (Section 5). Sample JU13 contains somewhat larger leucosome domains and <2 cm long cordierite-rich aggregates. This sample was split into a leucosome-poor (JU13A) and a leucosome-rich (JU13B) part. A coarse- to medium-grained red K-feldspar porphyritic granite (Romele granite) was sampled in a quarry at Veberöd (Fig. 3), where it occurs both as semiconcordant, cm-wide granitic layers in finely banded grey gneiss and as larger (>3 m) pods (Fig. 4c and e). The analysed sample (MGO080025) was taken in a quarry wall exposing a >10 m long body of greyish red undeformed isotropic coarse-grained granite. It has tartar twinned microcline megacrysts with conspicuous irregular aureoles or small inlets of fine-grained quartz, feldspar and myrmekite (Fig. 6e and f), plagioclase (generally saussuritized), and quartz. Fe–Mg phases are brown biotite, oxides, and pale brownish-green chlorite (Fig. 6e). Accessory phases are apatite, zircon and monazite, the latter phases commonly associated with oxides, and occasional muscovite. 6.1. Analytical method Approximately 0.3 kg solid rock sample was milled into a finegrained powder. A Wilfley water panning table was used to recover heavy mineral separates. Magnetic mineral fractions were removed with a hand magnet. Zircons were hand-picked under a stereomicroscope in alcohol and mounted on double-faced tape together with the 1065 Ma Geostandards zircon 91500 (Wiedenbeck et al., 1995). The mount was cast in epoxy and polished to expose the central parts of the crystals. Pre-analysis back scattered electron (BSE) imaging was used to select location of analytical spots using a Zeiss Supra 35-VP field emission scanning electron microscope (SEM) with a Robinson back scatter detector (at the Evolutionary Biology Centre, Uppsala University). The mount was coated with c. 30 nm of gold and U–Th–Pb analyses were done with a Cameca IMS 1280 high mass-resolution instrument (Secondary Ionisation Mass Spectrometry, SIMS), at the NORDSIM facility, Swedish Museum of Natural History in Stockholm. The analytical procedures followed Whitehouse et al. (1999) and Whitehouse and Kamber (2005). The instrument was operated with a spot size ≤20 ␮m. After analysis, location of analysed spots was reexamined with both BSE and Cathodoluminisence (CL) imaging. The post-analytical imaging was done using a Hitachi S-4300N electron microscope (at the Department of Geology, Lund University). The software Isoplot/Ex (Ludwig, 2003) was used for age calculations. All analytical data is presented in Table 7 and inverted U–Pb Tera Wasserburg diagrams in Fig. 11. Unless otherwise stated, all ages are reported with 2 errors, Concordia ages are reported with MSWD-values of both concordance and equivalence, and without decay constant errors (no significant difference in age is recorded with decay constant errors included). Data from misplaced analytical spots, i.e., spots that cross-cut textural domain boundaries or the crystal–epoxy interface, were excluded from age calculations. All excluded data are indicated in Table 7. 6.2. Zircon morphology and texture All samples are comparably rich in zircon. Grain sizes and the general morphological and textural characteristics of zircon are broadly similar in the different samples, including the coarsegrained red granite. They are mostly anhedral to subhedral, slightly discoloured greyish to distinctly brownish grains. Fractures are common, typically associated with discolouring, and most grains have abundant dark inclusions. The crystals are short prismatic with length-width aspect ratios below 2. They vary in size from <50 ␮m to >150 ␮m long. Smaller crystals (<30 ␮m) are typically

17

less fractured, less turbid, and have fewer visible inclusions than large crystals (>150 ␮m). BSE-images reveal considerable internal textural complexity (Fig. 10). BSE-dark, texturally old, core domains are oscillatory zoned, weakly zoned, or unzoned. These domains have been partly replaced by BSE-brighter, unzoned or weakly irregularly zoned, secondary domains. Some BSE-bright domains have a relict oscillatory zonation. Others are unzoned and crosscut the oscillatory zoning of the older domains; in some crystals they form distinct outer rims. In many crystals, the BSE-bright zircon forms deep embayments into BSE-dark zircon. The contact between texturally old and secondary domains is commonly distinct and locally follows fractures or primary zonation bands in the old oscillatory zoned zircon. In places it is associated with abundant inclusions or has a sponge like texture. The BSE-brighter secondary zircon domains are, as a rule, richer in U (often >500 ppm) than the texturally older domains (Table 7). Sections of smaller crystals (<50 ␮m across) are often unzoned, texturally non-complex, and BSE-bright (Fig. 10). The granite sample (MGO080025) contains a higher proportion of texturally non-complex crystal sections which are either BSE-bright, BSE-dark unzoned, or, more rarely, BSE-dark oscillatory zoned (Fig. 10, n3286-15a). The latter variety was only found in the granite sample. The results from the U–Th–Pb analyses are presented below. Textural classifications of the analysed points (domain location, zoning and BSE-intensity) and the isotopic data are presented in Table 7. Core–rim classification has only been applied in texturally complex zircon that shows a clear core–rim relation between the different zircon domains. If the exposed crystal section is texturally non-complex it is referred to as a non-complex section. This classification refers only to the section exposed in the polished section of the zircon mount and it does not imply that the whole grain is texturally non-complex. 6.3. Analytical results 6.3.1. JU12: Cordierite-garnet migmatitic granofels with weakly segregated leucosome Twelve analyses were made in zircon from sample JU12. The analyses define two distinct age groups at ca. 1.7 Ga and 1.45 Ga (Fig. 11a). The older group is defined by three concordant analyses in texturally old oscillatory zoned or slightly blurred primary core domains with relatively low-U contents and high-Th/U ratios of 0.2–0.8 (Table 7 and Figs. 10 and 12). The younger group is defined by nine analyses in texturally young, BSE-bright, unzoned or weakly zoned, thick (>30 ␮m) secondary replacements, rims or non-complex sections. They are distinctly higher in U (typically 600–900 ppm) and have lower Th/U-ratios <0.01–0.07 than the 1.7 Ga core domains (Fig. 12). Seven of these analyses are concordant and define a concordia age of 1448 ± 9 Ma (MSWD = 2.3, 95% conf.), identical to the weighted average 207 Pb/206 Pb age of 1447 ± 11 Ma (MSWD = 3.9, 95% conf.). The zircon in sample JU12 is thus complex and composed of igneous 1.7 Ga cores and extensive secondary zircon domains 1448 ± 9 Ma old. 6.3.2. JU13A: Leucosome-poor part of Cordierite-Garnet migmatitic granofels Eighteen analyses were made in zircon from the leucosomepoor split of sample JU13. Analogous to sample JU12, the analytical data plot in two distinct age clusters at 1.7 Ga and 1.45 Ga (Fig. 11b). Five spots were placed in texturally old BSE-dark core domains. One spot was placed too close to the crystal margin and was slightly reversely discordant (3%, analysis 17a). The remaining four spots define a common concordia age of 1699 ± 8 Ma (MSWD = 2.0; weighted average 207 Pb/206 Pb age of 1692 ± 9 Ma, MSWD = 1.1). Thirteen spots were placed in texturally young, BSE-bright unzoned or weakly irregularly zoned rims and non-complex crystal sections.

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Analysed area

Domain

Zoning

Comment

[U]

[Th]

[Pb]

Th/Ub

206

Pb/204 Pb

ppm

ppm

ppm

calc

measured

f206 Pbc

Ratios ± (%)

%

206

Pb/238 Pb

Ages ± (Ma) 207

Pb/206 Pb

0.12 {0.02} 0.03

0.3006 0.3068 0.3022

JU13A: Leucosome-poor part of cordierite-garnet migmatitic granofels BSE-bright, unzoned to irregularly zoned, rims or simple sections. Relatively enriched in U (>500 ppm) and with low Th/U ratios (<0.04) Simple wk irr 642 32 181 0.05 92759 0.02 0.2555 n4382-04a section Simple wk irr 691 53 197 0.07 152026 0.01 0.2561 n4382-05a section n4382-06a Rim uz 671 30 188 0.04 92781 0.02 0.2544 0.04 103423 0.02 0.2565 260 922 36 Simple uz n4382-08a section Simple uz 854 20 237 0.02 143566 0.01 0.2530 n4382-10a section n4382-13a Simple uz 546 52 153 0.09 94558 0.02 0.2511 section Rim wk irr 673 19 189 0.03 75642 0.02 0.2555 n4382-14a 0.03 125883 0.01 0.2584 330 36 1164 Simple uz n4382-01a section Simple uz 784 9 220 0.01 77940 0.02 0.2570 n4382-11a section Simple uz Straddles 592 31 174 0.05 98149 0.02 0.2659 n4382-02a section margin Simple uz Straddles 700 48 207 0.07 130476 0.01 0.2671 n4382-03a section margin n4382-12a Simple uz Straddles 847 39 248 0.05 48906 0.04 0.2655 section margin

206

Pb/238 Pb

0.25

1457

5

1463

13

0.40

1441

8

1455

13

0.35

1443

7

1454

14

0.31

1426

6

1449

13

0.45 0.44 0.26

1442 1447 1457

9 8 5

1421 1447 1452

13 13 13

0.35 0.28

1448 1449

7 5

1486 1483

13 13

0.2 0.2

0.98 1.01 0.96

0.1043 0.1072 0.1027

1.14 0.55 0.48

1702 1753 1673

21 10 9

1694 1725 1702

15 15 14

0.96

0.0916

0.47

1458

9

1467

13

1.02

0.0907

0.35

1441

7

1470

13

0.95 0.96

0.0911 0.0908

0.38 0.36

1450 1441

7 7

1461 1472

12 13

0.99

0.0910

0.32

1447

6

1454

13

0.95

0.0909

0.34

1444

7

1444

12

0.99 0.95

0.0912 0.0907

0.30 0.25

1451 1441

6 5

1467 1482

13 13

0.8

0.95

0.0905

0.30

1436

6

1474

13

0.5

0.96

0.0909

0.41

1443

8

1520

13

2.9

0.99

0.0908

0.34

1442

6

1526

13

3.8

1.03

0.0909

0.28

1444

5

1518

14

3.1

ARTICLE IN PRESS

BSE-dark oscillatory zoned (or convoluted zoning) cores. Relatively poor in U and with high Th/U ratios (>0.2) Core oz 45 24 17 0.50 14982 n4386-01a Core conv. 134 30 48 0.20 103252 n4386-02a 215 104 0.80 64502 Straddles 259 oz Core n4386-09a edge

Pb/206 Pb

J. Ulmius et al. / Precambrian Research xxx (2015) xxx–xxx

JU12: Cordierite-garnet migmatitic granofels with weakly developed leucosome patches BSE-grey to BSE-bright unzoned to weakly oscillatory zoned rims or simple sections. Relatively enriched in U (>c. 300 ppm) and with low Th/U ratios (<0.07) Simple uz 932 63 264 0.07 548371 {0.00} 0.2547 0.99 0.0915 n4386-03a section Simple uz 694 24 193 0.03 166991 0.01 0.2531 0.96 0.0907 n4386-04a section Simple uz 597 20 166 0.03 139797 0.01 0.2531 1.06 0.0908 n4386-05a section uz 21 177 0.03 148229 0.01 0.2520 0.98 0.0900 639 Simple n4386-06a section 76195 {0.02} 0.2466 0.98 0.0908 79 0.00 294 1 n4386-07a wk oz Rim Rim uz 608 25 169 0.04 181433 0.01 0.2517 1.00 0.0910 n4386-08a 0.2527 0.97 0.0915 0.01 0.06 175542 48 235 839 Simple uz n4386-11a section 0.0911 0.2593 1.00 0.01 250 0.05 238699 41 wk bands 872 Rim n4386-10a n4386-12a Rim wk oz 748 53 215 0.07 177437 0.01 0.2587 0.99 0.0911

207

Disc. % 2 limd

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Table 7 SIMS U–Th–Pb zircon data.

n4382-09a

Analysed area

Domain

Zoning

Simple section

uz

Comment

Ratios ± (%)

measured

%

206

242192

0.01

0.2472

0.95

0.0885

0.24

1394

5

1424

12

0.3012

0.95

0.1032

0.46

1683

8

1697

14

0.3007

0.97

0.1038

0.56

1694

10

1695

14

0.3036 0.3101

0.95 0.97

0.1043 0.1034

0.38 0.57

1701 1686

7 10

1709 1741

14 15

0.3

0.3166

0.96

0.1035

0.38

1688

7

1773

15

3.0

0.2543

0.96

0.0919

0.33

1465

6

1461

13

0.2551 0.2593

0.98 1.01

0.0919 0.0917

0.34 0.28

1466 1462

6 5

1465 1486

13 13

0.2607

0.96

0.0918

0.31

1464

6

1493

13

0.2720

1.02

0.0903

0.39

1431

7

1551

14

6.4

0.2676

1.03

0.0907

0.33

1440

6

1529

14

4.1

0.2664

0.96

0.0897

0.37

1418

7

1522

13

5.4

0.2599

0.99

0.0907

0.21

1439

4

1489

13

1.5

0.2491

0.95

0.0909

0.23

1444

4

1434

12

0.2308

1.12

0.0896

0.26

1417

5

1339

14

−3.6

0.1305 0.2099

1.36 1.70

0.0783 0.0879

0.68 0.35

1154 1381

13 7

791 1228

10 19

−28.9 −8.6

0.2539 0.2633

1.01 1.28

0.0907 0.0899

0.39 0.25

1440 1423

7 5

1459 1507

13 17

3.6

0.2106

1.16

0.0885

0.23

1393

4

1232

13

−10.3

0.1777

1.65

0.0851

0.23

1317

4

1055

16

−19

0.2373

1.56

0.0896

0.28

1417

5

1373

19

0

0.1201 0.1910

1.54 1.57

0.0771 0.0866

0.29 0.46

1124 1352

6 9

731 1127

11 16

−34 −15

[Th]

[Pb]

Th/Ub

206

ppm

ppm

ppm

calc

21

305

0.02

1129

Pb/204 Pb

BSE-dark, unzoned to weakly zoned (convoluted or broad banded) core domains relatively low in U and with high Th–U ratios (>0.6) Core wk conv Straddles 307 209 119 0.65 48275 0.04 n4382-07a rim Core wk conv Straddles 142 87 55 0.61 178255 {0.01} n4382-15a rim n4382-16a Core uz 303 231 122 0.77 >1e6 {0.00} Core uz Straddles 145 102 59 0.71 81337 {0.02} n4382-18a rim Core wk band Straddles 352 204 142 0.61 18590 0.10 n4382-17a margin JU13B: Leucosome-rich part of cordierite-garnet migmatitic granofels BSE-bright rims or simple sections of U-rich (>500 ppm) unzoned to weakly zoned zircon domains with low Th/U ratios (<0.11) Rim wk irr Straddles 742 36 209 0.05 35380 0.05 n4387-03a core n4387-04a Rim uz 672 70 192 0.10 102219 0.02 Rim wk irr Close to 765 13 217 0.02 154404 0.01 n4387-05a core n4387-07a Rim wk irr Straddles 636 36 184 0.06 62836 0.03 edge n4387-01a Simple uz Straddles 614 65 187 0.11 16427 0.11 section edge Rim wk oz Straddles 668 45 199 0.07 23398 0.08 n4387-02a edge Rim uz Straddles 547 39 161 0.07 10755 0.17 n4387-06a edge MGO080025: Red K-feldspar porphyritic near-isotropic granite (Romelegranite) BSE-bright replacement, recrystallised zircon, high-U (>1300 ppm), low Th/U (<0.1), unzoned or with relict oscillatory zonation Simple uz 1411 59 404 0.04 174289 0.01 n4384-01a section n4384-02a Simple uz Crack 1352 172 375 0.07 136418 0.01 section Eu rim uz Straddles 2031 71 515 0.03 39068 0.05 n4384-03a core Core uz Dark incl. 2561 91 364 0.02 17164 0.11 n4384-07a Rim uz Close to 1996 304 465 0.07 5008 0.37 n4384-08a margin Eu rim uz Thin crack 1298 77 365 0.05 517914 {0.00} n4384-09a Rim convolute Close to 1630 81 473 0.04 17361 0.11 n4384-10a margin Simple uz Crack 1878 647 440 0.07 27326 0.07 n4384-11a section n3286-02a Euhedral uz Close to 2055 132 406 0.07 43695 {0.04} rim crack Simple uz Cracked 1611 205 422 0.04 45462 {0.04} n3286-06a section Rim uz Cracked 2581 1583 353 0.13 4997 {0.37} n3286-14a Rim relict oz Cracked 1616 487 346 0.10 1664 {1.12} n3286-16a

Pb/238 Pb

Ages ± (Ma) 207

Pb/206 Pb

207

Pb/206 Pb

Disc. % 2 limd 206

Pb/238 Pb 0.1

ARTICLE IN PRESS

f206 Pbc

[U]

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Samplea / spot #

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Table 7 (Continued )

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Samplea / spot #

Analysed area

Domain

Zoning

Comment

[U]

[Th]

[Pb]

Th/Ub

206

Pb/204 Pb

ppm

ppm

ppm

calc

measured

BSE-dark, texturally old unzoned or oscillatory zoned core domains n4384-04a Core uz Close to 67 U-rich rim Core oz 190 n4384-12a 363 Cracked n3286-09a oz Core Core oz Cracked 546 n3286-10a Core wk oz Close to 476 n3286-11a crack Core wk oz Cracked 867 n3286-12a a b c d

Ratios ± (%)

%

206

{0.03} {0.95} {0.90} {1.05} {0.32}

0.2379 0.1773 0.1806 0.2232 0.2575

1.01 1.57 1.61 1.55 1.54

0.0905 0.0862 0.0836 0.0901 0.0862

1.32 1.89 3.72 1.54 2.18

1436 1342 1284 1428 1343

25 36 71 29 41

1376 1052 1070 1299 1477

12 15 16 18 20

{3.49}

0.1816

1.54

0.0836

3.99

1283

76

1076

15

{6.26}

0.2103

1.55

0.0878

3.99

1379

75

1231

17

{0.28}

0.2604

1.54

0.0906

0.77

1438

15

1492

21

{0.30}

0.2360

1.55

0.0892

1.31

1408

25

1366

19

Pb/238 Pb

Ages ± (Ma) 207

Pb/206 Pb

207

Pb/206 Pb

Disc. % 2 limd 206

Pb/238 Pb

51

25

0.57

28409

{0.07}

0.2872

0.95

0.1050

0.86

1714

16

1627

14

75 158 386 453

70 155 216 200

0.40 0.35 0.69 0.92

>1e6 6925 1628 27360

{0.00} {0.27} {1.15} {0.07}

0.3027 0.3514 0.3047 0.3076

0.97 1.54 1.54 1.54

0.1041 0.1187 0.1018 0.1034

0.50 0.33 0.68 0.43

1699 1937 1657 1686

9 6 13 8

1705 1941 1714 1729

14 26 23 23

221

184

0.27

454

{4.12}

0.1898

1.58

0.0640

1.38

743

29

1120

16

−13 −1

−1.3

41

Data used for age calculation shown with normal letters; data in italics have been excluded from age calculation. Th/U ratios calculated from 208 Pb/206 Pb ratios corrected for Pbcom. % of common 206 Pb in measured 206 Pb, estimated from 204 Pb assuming a present day Stacey and Kramers (1975) model. Degree of discordance; positive numbers are reverse discordant. Blanks indicate that analysis is concordant within 2 error. Abbreviations: conv, convoluted; irr, irregular; oz, oscillatory; wk, weak; uz, unzoned.

ARTICLE IN PRESS

BSE-dark newly crystallised unzoned or weakly oscillatory zoned zircon, low-U (majority <100 ppm), high-Th/U (>0.4) uz 53 47 17 0.83 63483 Core n4384-06a Core uz Cracked 59 52 14 0.69 1972 n3286-01a 17 11 4 0.42 2086 Core uz n3286-03a Core wk oz 81 92 25 0.83 1784 n3286-04a 1.21 5836 21 7 19 Simple uz n3286-05a section Simple uz 37 103 10 1.07 536 n3286-07a section Simple uz Cracked 35 73 12 1.39 299 n3286-13a section oz 146 114 50 0.81 6706 Simple n3286-15a section n3286-17a Eu core uz 47 47 15 1.02 6251

f206 Pbc

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Table 7 (Continued )

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Fig. 9. Calculated P–T pseudosections for the protolith compositions of cordierite-garnet-bearing granofels sample JU12 (a and b) and garnet-bearing biotite-sillimanite gneiss sample JU4 (c). Bulk compositions are given in Table 6. (a) Modal proportion contours (%) for melt (red) and (b) modal proportion contours (%) for garnet (red), sillimanite (blue) and staurolite (purple). Arrows represent suggested P–T path. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

Six analyses are discordant and excluded from age calculation. The discordant data comes from spots that were placed close to the crystal margin (analyses 02a, 03a and 12a), at a fracture (11a) or in domains with >1000 ppm U. The remaining seven concordant analyses define a concordia age of 1453 ± 6 Ma (MSWD = 1.0; weighted average 207 Pb/206 Pb age of 1447 ± 5 Ma, MSWD = 0.62). The secondary zircon domains have significantly higher U contents and lower Th/U ratios than the older domains (Table 7 and Fig. 12). In sample JU13A, zircon is made up of igneous domains dated at 1699 ± 8 Ma Ga and extensive secondary domains dated at 1453 ± 6 Ma.

6.3.3. JU13B: Leucosome-rich part of Cordierite-Garnet migmatitic granofels Seven analyses were made in zircon from the leucosome-rich part of sample JU13. All analyses cover texturally young secondary BSE-bright rims or texturally non-complex crystal sections (Fig. 10). Three analyses were misplaced and hit the crystal–epoxy interface; they yielded significantly reversely discordant data and were excluded from age calculation. The remaining four concordant analyses yield a concordia age of 1467 ± 7 Ma (Fig. 11c; MSWD = 1.1; weighted average 207 Pb/206 Pb age of 1464 ± 6 Ma, MSWD = 0.1). Texturally older BSE-darker zircon cores are

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common but too small for a ≥15 ␮m analytical spot to avoid mixing with surrounding younger high-U domains. The secondary formed zircon has high-U contents and low Th/U ratios analogous to the secondary zircon in the other samples (Fig. 12). In sample JU13B extensive formation of secondary zircon is dated at 1467 ± 7 Ma.

6.3.4. MGO080025: Sub-isotropic K-feldspar porphyritic red granite (Romele granite) Twenty-seven analyses were obtained but most analyses are distinctly discordant (Fig. 11d) and have a significant component of common Pb (206 Pb/204 Pb-ratios below 7000), which is anomalously low in comparison with the other samples in this study (cf. Table 7).

Fig. 10. BSE-images of representative zircon crystals showing crater after analytical spot and 207 Pb/206 Pb-age (Ma). (a) JU12, cordierite-garnet-bearing granofels. (b) JU13A, leucosome-poor part of cordierite-garnet-bearing granofels (c) JU13B, leucosome-rich part of cordierite-garnet-bearing granofels granofels. (d) MGO080025, K-feldspar porphyritic red Romele granite. Analytical ID refers to Table 7.

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Fig. 11. Tera Wasserburg diagram of U–Pb zircon SIMS analyses. Analyses used for age calculation are shown as filled error ellipses. (a) JU12, cordierite-garnet-bearing granofels. (b) JU13A, leucosome-poor part of cordierite-garnet-bearing granofels (c) JU13B, leucosome-rich part of cordierite-garnet-bearing granofels granofels. (d) MGO080025, K-feldspar porphyritic red Romele granite. Analytical ID refers to Table 7.

Post-analytical examination of the spot locations showed that the high common Pb-content in most cases associates with a spot location close to or across a fracture or the crystal–epoxy interface (commented in Table 7). Fractures are, however, not an explanation in the case of ca. 1.45 Ga BSE-dark, low-U cores and non-complex sections that have 206 Pb/204 Pb-ratios well below 7000. These analyses were made on a different mount in an earlier analytical session for which a significant component of common Pb was indicated for all zircon samples, and thus most likely an artefact introduced during preparation and handling of the mount (spot-number series n3286). Analogous to zircon in the cordierite-garnet-bearing granofels samples, the concordant data from the granite zircons plot in two groups at 1.70 Ga and 1.45 Ga. In addition, one concordant analysis of a BSE-dark oscillatory zoned texturally old core domain has a 207 Pb/206 Pb age of 1937 ± 12 Ma (Fig. 10, n3286-09). Three concordant spots in texturally old BSE-dark unzoned or oscillatory zoned domains define the older age group with a concordia age at 1691 ± 22 Ma (MSWD = 2.6; weighted average 207 Pb/206 Pb age of 1685 ± 44 Ma, MSWD = 3.6; Fig. 11). Twenty-one analyses define the younger ca. 1.45 Ga old age group. The young ages were

obtained from two distinctly different domains: (I) BSE-bright with high-U contents (> c. 1300 ppm) and low Th/U ratios, occurring as either non-complex (homogeneous) sections or as unzoned or relict oscillatory zoned rims/replacements, and (II) BSE-dark domains with low U contents (mostly <100 ppm) and high Th/U ratios (>0.4; Figs. 11 and 12), occurring either as non-complex sections or unzoned to oscillatory zoned cores. Most of the young analyses yielded highly discordant data and/or large analytical errors, and an age difference between these two types of younger zircon could not be discerned. Three concordant analyses with confined analytical errors (i.e., 207 Pb/206 Pb 2 error <40 Ma) define a common concordia age of 1445 ± 8 Ma (MSWD = 1.6; weighted average 207 Pb/206 Pb age of 1443 ± 7 Ma, MSWD = 0.18). These analyses represent a BSEbright euhedral rim (n4384-09), a BSE-bright non-complex section (n4384-02a) and a BSE-dark oscillatory zoned non-complex grain; cf. Fig. 10. To summarise, sample MGO080025 contains igneous zircon dated at 1691 ± 22 Ma [N = 3(5)], and extensive secondary zircon in the form of modified/recrystallised domains and new igneous crystals dated at ca. 1445 ± 8 Ma [N = 3(21)]. One inherited oscillatory zoned zircon core was dated at 1937 ± 12 Ma.

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Fig. 12. Th/U ratios versus 207 Pb/206 Pb age for U–Th–Pb SIMS analyses of zircon.

7. Discussion and conclusions 7.1. P–T path The investigated aluminous gneisses at Romeleåsen show that high temperatures prevailed during metamorphism. The evidence includes: • Sillimanite as the sole aluminosilicate polymorph. • Coexistence of cordierite + K-feldspar + sillimanite in granofels. • Abundant occurrence of leucosome, demonstrating partial melting. • Garnet is compositionally homogeneous (except along the outermost rims), likely homogenised at high temperature by intracrystalline diffusion. • Cordierite and garnet are present in leucosome, diagnostic of biotite dehydration melting. These findings show that temperatures peaked at or above 700 ◦ C. The peak metamorphic conditions for cordierite-garnetbearing granofels and garnet-bearing biotite-sillimanite gneiss at Romeleåsen are constrained at 700–750 ◦ C and 4–5 kbar (Fig. 8). By utilising pseudosections with melt reintegration, the early prograde history can be investigated. The subsolidus part of granofels JU12 as well as the gneiss JU4 is dominated by biotite, muscovite and aluminium silicate (Fig. 9). Garnet is restricted to pressures above 4–6 kbar and cordierite to pressures below 2–3 kbar. Staurolite is present between 3 and 8 kbar at 550–625 ◦ C, and is replaced by chlorite to lower temperatures. The petrographic observation of staurolite and sillimanite inclusions in garnet indicates that staurolite and sillimanite were present during prograde metamorphism. A prograde path that crosses the garnet-in line in the staurolite field gives the opportunity for staurolite inclusions in garnet and both garnet and staurolite will increase in modal content during the shown P–T path in the staurolite field (Fig. 9b). The discontinuous Reaction (1) can result in both reactant staurolite and product sillimanite as inclusions in growing garnet. At somewhat higher temperatures garnet would be introduced after sillimanite,

also rendering inclusion of sillimanite possible; however, this path would pass through the andalusite stability field. Pseudomorphs after andalusite have not been observed, and the presence of staurolite inclusions is not compatible with such a path. The difference between the staurolite field and the peak metamorphic conditions is 100–150 ◦ C at about the same pressure, implying a prograde path dominated by heating. Formation of cordierite at peak metamorphic conditions reflects crossing of the melting reaction involving the consumption of biotite and sillimanite to produce cordierite, garnet, K-feldspar and melt. At the biotite-out line, the modal content of cordierite reaches 10–15 mol%, which is observed for granofels of Romeleåsen. The modelling predicts that with a temperature higher than 750 ◦ C biotite would disappear completely. This is not the case as sparse grains of reddish-brown biotite are present in the cordieritegarnet-bearing granofels. As already mentioned, the stability field of biotite may be enlarged to higher temperatures due to the high Ti-content in biotite, which is not reproduced in the models. The formation of biotite and sillimanite at the expense of garnet, and the formation of pseudomorphs after cordierite with biotitesillimanite-muscovite, points to continuous retrogression below 650 ◦ C, with supply of fluids, likely at simultaneously decreasing pressure. A clockwise P–T path is suggested as: • Inclusions of staurolite occur in garnet. • A counter-clockwise or isobaric heating path across the low-P part of the staurolite field would pass through the andalusite field and/or fields without garnet. The proposed path in Fig. 13 indicates that prograde heating was associated with limited burial, whereas limited exhumation occurred during retrograde cooling. This path is in agreement a clockwise evolution (cf. Thompson and England, 1984), but at significantly higher temperatures and lower pressures than typical for Barrovian metamorphism. The P–T evolution is characteristic of accretionary settings, as e.g., the Lachlan orogeny in eastern Australia (Coney, 1992; Collins, 2002a,b). The metamorphic peak at 4 kbar and 700 ◦ C for the Romeleåsen rocks corresponds to an average geothermal gradient of 50 ◦ C/km (assuming 1 kbar = 3.5 km).

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(available age data summarised and referenced in Fig. 2 and Table 1). The data supports models for extensive ductile deformation, high-grade metamorphism and crustal anatexis in Baltica at 1.47–1.38 Ga (Hallandian orogeny). The relict high-grade Hallandian gneiss belt inside the Eastern Segment of the Sveconorwegian orogen is cut, at high-angle, at the Mylonite Zone (Fig. 2). Corresponding 1.47–1.38 Ga high-grade metamorphism and deformation has not recognised in the western Sveconorwegian terranes (cf. Fig. 2). 7.3. 1.47–1.44 Ga magmatism in Baltica

Fig. 13. Inferred P–T path from petrographic textures and pseudosection calculations for aluminous rocks from Romeleåsen.

7.2. 1.47–1.38 Ga metamorphism in Baltica Texturally young, U-rich (>500 ppm), BSE-bright, secondary zircon with low Th/U ratio (<0.1), is present in all investigated samples and is dated at 1.45 Ga. It occurs as non-complex sections, texturally young distinct rims, transitional replacements or as deep embayments into older cores. The coherent isotopic and textural character of the secondary zircon in all samples suggests that it formed at the same time and by a similar process. Zircon is a mineral highly robust to reworking, also at elevated temperatures, and the U–Pb isotopic system in crystalline zircon with normal U-content is a closed system at temperatures <800 ◦ C (Cherniak and Watson, 2003). The presence of melt may, however, catalyst in situ recrystallisation of zircon (coupled dissolution–reprecipitation), or alternatively, zircon may dissolve in zircon undersaturated melt, and later precipitate as new grains and overgrowths (Hanchar and Watson, 2003; Harley et al., 2007). All samples investigated here are associated with melt, either as in situ partial melt patches (migmatitic granofels samples JU12 and JU13) or as melt assembled to form layers and small bodies of granite (granite sample MGO080025). The petrography and thermodynamic modelling suggests that peak temperatures for metamorphism did not exceed 750 ◦ C (Section 7.1 above). It is therefore suggested that the reworking and new formation of zircon was associated with reactions between zircon and melt. The BSE-bright, U-rich secondary zircon in the three granofels samples is dated at 1467 ± 7 Ma, 1448 ± 9 Ma and 1447 ± 5 Ma, respectively, and directly date the reaction between pre-existing zircon and melt. The lack of a statistical overlap between the older age of 1467 ± 7 Ma (N = 4 in JU13B) and the younger ages (JU12 + JU13A) is unclear. The statistical basis (N = 14) for the younger ages is better and a pooled age of these analyses in secondary zircon (JU12 + JU13A) is calculated as a concordia age of 1451 ± 6 Ma (MSWD = 1.6) which directly sets the age for Hallandian high-grade metamorphism. The 1451 ± 6 Ma age for high-grade metamorphism, including partial melting at Romeleåsen is analogous to ages for early gneissic layering and partial melting in the Eastern Segment of the Sveconorwegian province and the Blekinge-Bornholm province

Zircon in the unmetamorphosed and undeformed red porphyritic granite (Romele granite) contains ca. 1.45 Ga BSE-dark newly crystallised oscillatory zoned or unzoned zircon with low Ucontent and high-Th/U-ratio; a type absent in the granofelses (cp. Fig. 12). The age of this oscillatory zoned zircon does not differ from the age of the BSE-bright zircon in the same sample and a common age of 1445 ± 8 Ma dates the igneous crystallisation of the Romele granite. The 1.45 Ga zircon data from the migmatitic granofels and the Romele granite shows that the granite formed at the same time as crustal anatexis of the surrounding gneisses. The 1445 ± 8 Ma igneous crystallisation age of the Romele granite supports previous suggestions by Hjelmqvist (1934) that red K-feldspar porphyritic granites at the Romeleåsen horst are related to granites in the Blekinge-Bornholm province (Åberg, 1988) and that these belong to the 1.47–1.44 Ga granitic to quartzmonzodioritic intrusions in the southern Baltic Shield (Fig. 2; cp. Brander and Söderlund, 2009). Concomitant 1455 ± 6 Ma massif type anorthosites are preserved in the eastern marginal parts of the Sveconorwegian orogen (Jönköping Anorthosite Suite, Brander and Söderlund, 2009). The 1.47–1.44 Ga magmatism appears to increase in volume towards the south, with abundant and large intrusions in the Blekinge-Bornholm province (Fig. 2; Kornfält and ˇ cys Vaasjoki, 1999; Geisler and Schleicher, 2000; Åhäll, 2001; Ceˇ et al., 2002; Söderlund et al., 2002; Obst et al., 2004; Möller et al., 2007; Zarin¸sˇ and Johansson, 2009; Petersson et al., 2013; Waight et al., 2012). 1.46–1.44 Ga old anorthosite-mangerite-charnockitegranite associations have also been documented in the subsurface East European Craton (Motuza et al., 2006; Skridlaite et al., 2007). Further into the Baltic Shield (ca. 500 km north of Blekinge) 1.46 Ga mafic dykes and sills form a ca.1000 km traverse across from the Caledonian front in the west to Lake Ladoga in the east (Fig. 1; Söderlund et al., 2005; Brander and Söderlund, 2009; Lubnina et al., 2010). Records of 1.47–1.44 Ga magmatism are conspicuously rare in the western terranes of the Sveconorwegian orogen (west of the Mylonite Zone); to our knowledge it has been recorded at two places only [Islandsberg dyke north (Orust dykes) north of Orust in the Idefjorden terrane (Åhäll and Connelly, 1998), and banded gneiss in the Telemarkia terrane of southern Norway (Pedersen et al., 2009)]. In addition, a partly well preserved mafic dyke swarm in the Idefjorden terrane (Koster dyke swarm) has given a Rb–Sr whole rock age of 1421 ± 25 Ma (Hageskov and Pedersen, 1988). 1.47–1.44 Ga old gabbroic intrusions are also known from the central Norwegian Caledonides (Corfu and Emmett, 1992; Corfu and Andersen, 2002; Tucker et al., 2004; Krogh et al., 2011; Beckman et al., 2014). Today, the Hallandian orogenic province is cut at high angle by the younger Sveconorwegian orogen which in turn is cut by the Caledonides (Fig. 1). Both orogenies caused extensive reworking of crust and tectonic displacement hundreds of kilometres into Baltica. The original extent of Hallandian magmatism in western and southern parts of Baltica is unknown; it may have been far more extensive than the remnants exposed today in the southernmost shield area. The scarcity of Hallandian 1.47–1.38 Ga magmatism

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in the western Sveconorwegian terranes, further emphasise the role of the Mylonite Zone as a first-order lithotectonic collisional boundary of Sveconorwegian age. 7.4. Geological interpretation of composite zircon in polymetamorphic gneisses: potential pitfalls Zircon in the undeformed and unmetamorphosed Romele granite (sample MGO080025), contains abundant xenocrystic 1.7 Ga cores surrounded by 1.45 Ga U-rich, BSE-bright envelopes. The high-U zircon dates the emplacement of the unmetamorphosed granite at about 1.45 Ga (Sections 7.2–7.3 above). Identical zircon poulations, with BSE-dark oscillatory zoned 1.7 Ga cores surrounded by ca. 1.45 Ga U-rich, BSE-bright envelopes are typically found in polymetamorphic high-grade orthogneisses in the Eastern Segment (Fig. 2; Table 1; Söderlund et al., 2002; Möller et al., 2007; Rimˇsa et al., 2007; Brander et al., 2012). In these rocks, however, the 1.7 Ga cores are interpreted to date igneous crystallisation of the orthogneiss protolith, while the 1.45 Ga secondary zircon envelopes are interpreted to date migmatisation of the gneiss. In most cases, this interpretation is based on the fact that the volume of the secondary zircon formed increase with increasing volume of leucosome, whereas secondary zircon is incapacious to almost absent in unveined or sparsely veined orthogneiss (e.g., Söderlund et al., 2002; Möller et al., 2007, cp. also Petersson et al., 2013). A different and tentative interpretation was presented for unveined meta-intrusions, described as weakly gneissic only, in the eastern marginal parts of the Eastern Segment (Table 1; Brander et al., 2012). The formation of thick, ca. 1.45 Ga, BSE-bright, Urich rims on 1.70 Ga BSE-dark cores, identical to those in the Romele granite, was interpreted to reflect solid state metamorphism (amphibolite-facies) of a 1.70 Ga granite (Brander et al., 2012). In this case there is no evidence as to whether or not the sub-solidus amphibolite-facies metamorphism could have caused the recrystallisation of zircon; it is actually less probable in the light of reported zircon systematics for unveined orthogneisses in the Eastern Segment (Söderlund et al., 2002, Möller et al., 2007, Petersson et al., 2013). An alternative interpretation is that, similar to the Romele granite, the 1.45 Ga zircon formed during emplacement and crystallisation of the granite itself, and that 1.7 Ga zircon is xenocrystic. If so, the extent of 1.45 Ga granite magmatism may be strongly underestimated. We emphasise that the interpretation of composite zircon is dubious if a connection between a geological process and the formation of zircon is missing. For example, if a granite with an initially complex zircon population is overprinted by metamorphism and zircon stayed inert, the youngest zircon phase could easily be mistaken to date metamorphism of an older granite intrusion, while in fact, the youngest zircon phase dates the igneous crystallisation of the granite, not metamorphic crystallisation. 7.5. Interpretation of age and origin of aluminous gneisses 7.5.1. Origin of aluminous gneisses Aluminous gneisses are commonly interpreted to be of sedimentary (epiclastic) origin, but chemical weathering or hydrothermal alteration may cause substantial relative enrichment of Al (and Si) in an igneous rock (Schwartz, 1939; Middelburg et al., 1988). Subsequent penetrative metamorphic recrystallisation and ductile deformation may obliterate field relations and the chemical, structural and textural character of the protolith may become completely obscured. Such process will form Al-silicate-bearing gneisses of a supracrustal but not necessarily a sedimentary (epiclastic) origin.

Occurrences of aluminous metamorphic rocks are known from the Eastern Segment and the Blekinge Bornholm province (Fig. 1; Larsson, 2001). They formed by acid leaching of granitic and volcanic rocks in a hydrothermal magmatic environment at shallow to supracrustal levels (Larsson, 2001). The metasomatic alteration caused relative enrichment of Al and Si, and subsequent metamorphism formed Al-silicate. Similar high-alumina deposits are also known from other parts of the Baltic Shield (Fig. 1; Hallberg and Fallick, 1994; Poutiainen and Grönholm, 1996). Furthermore, in the Palaeoproterozoic volcanic province of Bergslagen, syn-volcanic hydrothermal alteration of igneous rocks is common and metamorphic equivalents of these rocks have, based on the presence of Al-silicate, in many places been misinterpreted as epiclastic deposits (Stephens et al., 2009). The aluminous gneisses at the Romeleåsen horst may be of a composite volcanic and sedimentary origin, as suggested by Hjelmqvist (1934). Volcanic associations are common in the Transscandinavian Igneous Belt (cf. Andersson and Wikström, 2004; Appelquist et al., 2008), but epiclastic sediments are rare. It is possible that the high-alumina rocks at Romeleåsen have been hydrothermally altered, analogous to aluminous rocks elsewhere in the Eastern Segment (Fig. 1; cf. Lundegårdh, 1995; Larsson, 2001), however to a lesser extent as they show mild depletion only in alkali elements. The relative enrichment of Al and Si in these rocks could also have resulted from chemical weathering and subsequent removal of dissolved elements by infiltrating fluids. Both scenarios imply a supracrustal, but not necessarily an epiclastic, origin. An epiclastic origin was recently proposed (Bogdanova et al., 2014) for aluminous migmatitic gneisses exposed at another northwest-trending Precambrian basement horst ca. 40 km northeast of Romeleåsen (Linderödsåsen, Fig. 2 and ID number 3–5 in Table 1). The investigated rocks are described as migmatitic garnet-biotite-bearing gneisses with cordierite, sillimanite and/or muscovite of a possible wacky to arkosic origin. However, these rocks are similar to the aluminous gneisses described in this study; the metamorphic P–T-conditions were estimated as similar to those at Romeleåsen, and it is suggested that also these aluminous gneisses may have formed by chemical weathering or hydrothermal alteration of an igneous protolith (cf. also Section 7.5.2 below). 7.5.2. The protolith age of the aluminous gneisses All zircon samples investigated in this study contain relics of 1.70 Ga BSE-dark, unzoned or oscillatory zoned, zircon cores (typically fractured) with comparably low U-content (<400 ppm) and high Th/U-ratio (>0.2). One zircon core from the granite sample is dated at 1.94 Ga. The common textural and isotopic characteristics of the protolith zircon suggest a common igneous, almost exclusively 1.70 Ga old, source. 1.70 Ga old orthogneiss protolith ages have previously been demonstrated at the Romeleåsen horst, (red aplitic gneiss at Stenberget, Johansson et al., 1993) and dominates among orthogneisses in the Eastern Segment of the Sveconorwegian province (Connelly et al., 1996; Christoffel et al., 1999; Söderlund et al., 1999, 2002; Andersson et al., 2002; Möller et al., 2007; Brander et al., 2012; Petersson et al., 2013). These rocks are reworked equivalents to the younger 1.74–1.66 Ga magmatic rocks of the Transscandinavian Igneous Belt (Figs. 1 and 2; TIB 2 and 3; op. cit; Andersson and Wikström, 2004). In this study, the relict 1.70 Ga zircon core domains were difficult to date due to extensive fracturing of the cores and wide-spread occurrence of secondary U-rich domains, often penetrating deep into the grains. Such problems are common for analyses of complex zircon with both high- and low-U phases because (I) volume expansion of U-rich domains may cause fracturing of the low-U domain (Corfu et al., 2003), and (II) the high-U phase will strongly influence the U–Pb isotopic system even at small degrees of mixing. Reliable

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ages for U-poorer fractured relict cores surrounded by extensive U-rich envelopes, such as the complex zircon investigated in this study, therefore difficult to obtain since incorporation of even tiny amounts of young U-rich zircon will yield erroneously young ages, and fractured domains typically host common-Pb resulting in substantial correction and high analytical uncertainties. U–Pb zircon data from complex zircon with extensive secondary U-rich envelopes around fractured relict cores poorer in U was recently reported from aluminous gneisses in the BlekingeBornholm province (Fig. 2; ID-number 3–5 in Table 1; Bogdanova et al., 2014). The data show a significant spread in U–Pb ages, between 1.70 and 1.50 Ga, and was interpreted to reflect different igneous ages of detrital zircon grains (Bogdanova et al., 2014). An alternative explanation is that the spread of ages and discordance of data points result from mixing between 1.70 Ga fractured protolith cores (possibly hosting common Pb) and 1.45 Ga secondary high-U zircon envelopes characteristic for the analysed zircon populations (see above). Several examples of analogous spread of ages caused by mixing of high-U secondary zircon and fractured relict cores in high-grade orthogneiss have been published from other parts of the southern Baltic Shield (e.g., in Möller et al., 2007, Rimˇsa et al., 2007, and Brander et al., 2012). If the spread of ages is caused by mixing the data will form an array in plots of age versus Th/U or U. The Romele granite and the aluminous gneisses dated by Bogdanova et al. (2014) also contain a minor component of 1.94 Ga igneous zircon. Igneous 1.94 Ga rocks are absent in the southern shield areas and, except for in the Lappland Kola-belt in the northeasternmost parts, rare also in the northern shield (e.g., Wasström, 2005; Bogdanova et al., 2006). Nevertheless, the 1.94 Ga component is present in both investigated exposures in the southernmost Baltic Shield (the Romele and the Linderöd horsts; Bogdanova et al., 2014; this study). Another noteworthy feature is the absence of 1.82–1.77 Ga components, which is the age of the presently exposed older bedrock in the Blekinge-Bornholm province, immediately northeast thereof (Fig. 1; cf. Johansson et al., 2006). 7.6. The Hallandian orogeny: setting and context In this paper we present the first evidence for a Hallandian orogenic clockwise P–T path under a strongly elevated geotherm, coeval with partial melting and granite intrusion at 1.45 Ga. Together with the regional data this shows that Hallandian orogeny in Baltica caused regional crustal-scale high-temperature metamorphism and partial melting at low to moderate pressures (4–5 kbar) and was accompanied by ductile deformation. The high-temperature metamorphism was concomitant with granitic magmatism, also producing subordinate amounts of granodiorite, ˇ cys and Benn, 2007; Zarin¸sˇ and Johansson, quartz-monzonite (Ceˇ 2009), and anorthosite (Brander and Söderlund, 2009). Magmatism was extensive in the parts of the Hallandian orogen exposed in the Blekinge-Bornholm province. Inside the Eastern Segment of the Sveconorwegian orogen to the west, Hallandian high-grade metamorphism involved migmatisation at 1.47–1.42 Ga (Table 1). The time of Hallandian magmatism (well-constrained published U–Pb ages) is in the same interval (Fig. 1; Söderlund et al., 2002; Möller et al., 2007; Brander and Söderlund, 2009; Zarin¸sˇ & Johansson, 2009; Waight et al., 2012). Deposition of sediments (sandstone, conglomerates, siltstones and shales) and injections of basaltic sills and dolerites took place further inward the craton (Fig. 1; Jotnian sediments, basalts and dolerites; Söderlund et al., 2005; Lubnina et al., 2010; Ripa et al., 2012; Gee et al., 2014). These ca. 1.46 Ga old mafic intrusions and the sedimentary basins in the foreland of the Hallandian orogen are exposed from Dalarna in the west to the Lake Ladoga region in the east (>900 km) and substantial parts of the offshore Jotnian sedimentary rocks and mafic rocks may also belong to this succession (Fig. 1). Based on airborne magnetic anomaly

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patterns (e.g., Bogdanova et al., 2015) and field surveying east of the Sveconorwegian Front, the structural grain of the Hallandian orogen appears to essentially follow the E–W to NW–SE-structural trend of the earlier Paleoproterozoic structures. The present day exposure of the Hallandian orogen in the Baltic Shield includes no known allochthonous boundary separating crustal components of different lithologic association, age, structure or metamorphism. High-pressure metamorphism is lacking. This is in striking contrast to the 1.1–0.9 Ga Sveconorwegian orogen which comprises crustal-scale continental blocks displaced out of their original context, as well as deeply depressed, highpressure metamorphic parts of Baltica (Möller et al., 2015). The Hallandian, 1.47–1.42 Ga, crustal-scale high-temperature, moderate–low-pressure metamorphism, ductile deformation and concomitant magmatism of the southern Baltic Shield and subsurface equivalents in the East European Craton is interpreted to reflect an active continental margin setting of Andean-type. Concomitant mafic magmatism and basin formation deeper into the foreland is interpreted to reflect back arc extension in response to subduction along the continental margin. Hallandian metamorphism is characterised by low to moderate pressures and a high-geothermal gradient. Available geochronological data points to an orogenic cycle with most of the magmatism bracketed in the 1.47–1.44 Ga interval. The high-temperature, moderate-to low pressure metamorphism and the associated wide-spread magmatism that characterise the remnants of Hallandian orogeny suggests an accretionary orogenic setting while evidence for a terminal phase of continent-continent collision, such as records of highpressure metamorphism of deeply buried or subducted continental crust of the Precambrian continent Baltica at 1.5–1.4 Ga, is lacking. Records of late- to post-orogenic Hallandian igneous activity are exposed in the southwestern coastal parts of the Eastern Segment. It includes dominantly granite-mangerite-norite-charnockite magmatism and associated, locally fluid-assisted, metamorphism at 1.42–1.38 Ga (Hubbard, 1975; Åhäll et al., 1997; Christoffel et al., 1999; Andersson et al., 1999; Rimˇsa et al., 2007; Harlov et al., 2006, 2013; Johansson et al., 2013). The character of this magmatism points to ponding of mafic magmas at the base of the crust, and could possibly indicate late-Hallandian delamination (cp. McLelland et al., 2010). The outline of the Hallandian orogen has been severely modified by substantial crustal-scale tectonic reconfigurations of the plate margin after the Hallandian orogeny (younger orogenic events and periods of continental rifting). The original tectonic build-up and extent of the Hallandian orogen and, importantly, its spatial relation to an active continental margin (the palaeo-margin of Baltica), is therefore concealed. However, our data suggest that an active continental margin can be expected at 1.47–1.38 Ga SW of Baltica, corroborating recent palaeogeographic reconstruction by Pisarevsky et al. (2014). 8. Conclusions Detailed petrology, thermodynamic P–T modelling and dating using U–Pb SIMS analysis of zircon of aluminous gneisses and spatially associated granites in the southeasternmost marginal part of the Sveconorwegian orogeny show that • Hallandian metamorphism records a clockwise orogenic P–T evolution reaching granulite-facies temperatures (700–750 ◦ C) at low pressures (4–5 kbar). • Hallandian high-grade metamorphism is dated at 1451 ± 6 Ma. • Hallandian regional metamorphism took place under a strongly elevated geotherm and was associated with granitic magmatism, characteristic of an accretionary orogenic setting.

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• Our data suggest an active continental (Andean-type) margin at the SW margin of Baltica for the period 1.47–1.38 Ga, placing constraints for paleogeographic reconstruction. • A common source is suggested for the zircons of the aluminous gneisses and the spatially associated granites with a protolith age of 1.70 Ga, which corresponds to the young generation of Transscandinavian Igneous Belt rocks. The aluminous gneisses may be of supracrustal, but not epiclastic, origin. Acknowledgements This project was financially supported by the Royal Physiographic Society in Lund (grant to C. Möller) and the Geological Survey of Sweden (geochronology program). Many thanks to Hugo Wikman and Karl-Axel Kornfält for geological guiding on the Romeleåsen horst and discussions on the geology. Mattias Göransson, SGU, is thanked for showing localities of the Romele granite and for sharing samples and thin sections. Carl-Henric Wahlgren is thanked for fruitful discussions on the regional geology of the southern Baltic Shield. Martin Whitehouse and the staff at the NORDSIM laboratory in Stockholm are thanked for technical guidance. The NORDSIM facility is operated under a contract between the research funding agencies of Denmark, Iceland, Norway, and Sweden, and the Geological Survey of Finland and Swedish Museum of Natural History. This paper is NORDSIM contribution No. 400. We also thank the editors Giulio Viola and Randall Parrish and reviewer John Diener and an anonymous reviewer for their constructive reviews and for improving the manuscript. References Åberg, G., 1988. Middle Proterozoic anorogenic magmatism in Sweden and worldwide. Lithos 21, 279–289. Åhäll, K.-I., 2001. Åldersbestämning av svårdaterade bergarter i sydöstra Sverige. SKB Report R-01-60, Stockholm. Åhäll, K.-I., Connelly, J., 1998. Intermittent 1.53–1.13 Ga magmatism in western Baltica; age constraints and correlations within a postulated supercontinent. Precambrian Res. 92, 1–20. Åhäll, K.-I., Connelly, J., 2008. Long-term convergence along SW Fennoscandia: 330 m.y. of Proterozoic crustal growth. Precambrian Res. 163, 402–421. Åhäll, K.I., Samuelsson, L., Persson, P.O., 1997. Geochronology and structural setting of the 1.38 Ga Torpa granite; implications for charnockite formation in SW Sweden. Geol. Fören. Stockh. Förh. 119, 37–43. Andersen, T., Andersson, U.B., Graham, S., Åberg, G., Simonsen, S.L., 2009. Granitic magmatism by melting of juvenile continental crust: new constraints on the source of Palaeoproterozoic granitoids in Fennoscandia from Hf isotopes in zircon. J. Geol. Soc. Lond. 166, 233–247. Andersson, U.B., Wikström, A., 2004. The Småland-Värmland Belt. In: Högdahl, K., Andersson, U.B., Eklund, O. (Eds.), The Transscandinavian Igneous Belt (TIB) in Sweden: a review of its character and evolution. Geol. Surv. Finland Spec. Pap. 37, 15–20. Andersson, J., Söderlund, U., Cornell, D., Johansson, L., Möller, C., 1999. Sveconorwegian (Grenvillian) deformation, metamorphism and leucosome formation in SW Sweden, SW Baltic Shield: constraints from a Mesoproterozoic granite intrusion. Precambrian Res. 98, 151–171. Andersson, J., Möller, C., Johansson, L., 2002. Zircon geochronology of migmatite gneisses along the Mylonite Zone (S Sweden): a major Sveconorwegian terrane boundary in the Baltic Shield. Precambrian Res. 114, 121–147. Andersson, U.B., Rutanen, H., Johansson, Å., Mansfeld, J., Rimsa, A., 2007. Characterization of the Paleoproterozoic Mantle beneath the Fennoscandian Shield: geochemistry and isotope geology (Nd, Sr) of ∼1.8 Ga mafic plutonic rocks from the Transscandinavian Igneous Belt in Southeast Sweden. Int. Geol. Rev. 49, 587–625. Appelquist, K., Cornell, D., Brander, L., 2008. Age, tectonic setting and petrogenesis of the Habo Volcanic Suite: evidence for an active continental margin setting for the Transscandinavian Igneous Belt. GFF 130, 123–138. Austin Hegardt, E., Cornell, D., Claesson, L., Simakov, S., Stein, H., Hannah, J., 2005. Eclogites in the central part of the Sveconorwegian Eastern Segment of the Baltic Shield: support for an extensive eclogite terrane. GFF 127, 221–232. Beckman, V., Möller, C., Söderlund, U., Corfu, F., Pallon, J., Chamberlain, K.R., 2014. Metamorphic Zircon Formation at the Transition from Gabbro to Eclogite in Trollheimen–Surnadalen, Norwegian Caledonides. Geol. Soc. Lond. Spec. Publ. 390, 403–424. Bergerat, F., Angelier, J., Andreasson, P.-G., 2007. Evolution of paleostress fields and brittle deformation of the Tornquist Zone in Scania (Sweden) during PermoMesozoic and Cenozoic times. Tectonophysics 444, 93–110.

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Please cite this article in press as: Ulmius, J., et al., Hallandian 1.45 Ga high-temperature metamorphism in Baltica: P–T evolution and SIMS U–Pb zircon ages of aluminous gneisses, SW Sweden. Precambrian Res. (2015), http://dx.doi.org/10.1016/j.precamres.2015.04.004