Proterozoic Vammala Migmatite Belt, Southern Finland, and Implications for Svecofennian Tectonics

Bulletin of the Geological Society of Finland,Vol. 76, 2004, pp. 93–140

Pre-1.91 Ga deformation and metamorphism in the Palaeo- proterozoic Vammala Migmatite Belt, southern Finland, and implications for Svecofennian tectonics

R. W. Roye Rutland1)*, Ian S. Williams1), and Kalevi Korsman2) 1) Research School of Earth Sciences, Australian National University, Canberra, ACT 0200, Australia 2) Pietiläntie 6 A2, 03100 Nummela, Finland

Abstract A metamorphic event in the Vammala Migmatite Belt (VMB) at ~1.92 Ga, revealed by SHRIMP U-Pb analyses of both zircon overgrowths and monazite,is interpreted as post-depositional and is correlated with the development of the early high-grade schistosity.Neither this Early Svecofennian deformation and metamorphism, nor the associated complex folding, is present in the overlying Ta mpere Schist Belt (TSB) sequence, consistent with the VMB being part of a pre-1.91 Ga basement complex.The ~1.92 Ga event provides a maximum deposition age for the TSB,confirming earlier age estimates. Earlier stratigraphic correlations between parts of the VMB and TSB,and associated tectonic interpretations, can no longer be sustained.The crustal thickening seen in the VMB,and previously attributed to arc accretion at ~1.89 Ga, is now attributed to accretion of a large Svionian marginal basin during the ‘Early Svecofennian’ orogenic phase at ~1.92 Ga.This is of similar age to the deformation and metamorphism associated with collision in the Lapland-Kola Orogen to the north of the Karelian Province. The well-known post-TSB orogenic phase was also identified in the VMB by a monazite age of 1881±6 Ma.A granitoid intrusion gave an emplacement age of 1888±5 Ma, comparable to the age of granitoid clasts in the upper part of the TSB succession. The detrital zircon data are interpreted to suggest that deposition of the precursor VMB sediments probably took place soon after an earlier pre-depositional metamorphism at ~1.98 Ga, which affected igneous source complexes dated at ~1.99 Ga and ~2.01 Ga. Mafic rocks in the southern part of the VMB,and probably also the Ha- veri basalts, represent a renewed episode of extensional magmatism, which might correlate with the 1.96–1.95 Ga Jormua and Outokumpu ophiolites. A pre-1.96 Ga older stage basin has an expression in Sweden and complexes of similar age occur in the concealed Palaeoproterozoic basement south of the Gulf of Finland. Similar rocks, deformed and metamorphosed before ~1.96 Ga, might be present beneath the Central Finland granitoid complex and the late Svecofennian granite-migmatite zone,and were possibly more local sources for both the younger stage Svionian basin sediments and the post-1.91 Ga Bothnian Basin sediments. The TSB and other post-accretionary volcanic sequences, and the associated plu- tonism, are interpreted to reflect a ~40 m.y.extensional period, inboard of the contemporaneous active margin, between orogenic phases at ~1.92 Ga and ~1.88 Ga.This interpretation provides a more satisfactory explanation of the major heat input to the crust over a very wide area than does the arc accretion hypothesis.The tectonic evolution of the Svecofennian Province has strong similarities to that of the Palaeozoic Lach- lan Fold Belt in eastern Australia.

Key words: metasedimentary rocks, migmatites, gneisses, deformation, metamorphism, tectonics, Svecofennian Domain, absolute age,U/Th/Pb,zircon, monazite,Palaeoprot- erozoic,Vammala,Ta mpere,Finland

* Corresponding author e-mail: [email protected] 94 Rutland,R.W.R.,Williams,I.S. and Korsman,K.

1. Introduction

The Svecofennian Province of the Fennoscandi- 1.87 Ga)”(A and BinFig.1), both of which an Shield is exposed in an approximately trian- have been considered to contain exotic terranes, gular area about 1,000 km on a side adjacent to mainly on the basis of geochemical and detrital the Archaean Karelian Province (Fig. 1). Similar zircon data (e.g. Lahtinen, 1994; Rämö et al., Palaeoproterozoic rocks arealso preservedbe- 2001; Lahtinen et al., 2002). The migmatites in neath the Phanerozoic cover of the East Europe- the north of the VMB have been interpreted as a an Platform (Gorbatschev & Bogdanova, 1993). fore arc (Lahtinen 1996, p. 43), or accretionary East of the post-1.85 Ga igneous complexes on complex (Kähkönen, 1999, pp. 26–27) closely its western margin, the Svecofennian Province related in time to the TSB, which was inferred proper is dominated by intrusive granitoids as- to have been formed above a northward dipping sociated with more limited volcanic sequences subduction zone. Lahtinen (1994, p. 95) and and extensive migmatite complexes of which the Kähkönen et al. (1994, pp. 44–45, Fig.1) also Vammala Migmatite Belt (VMB) is one of the suggested the presence of a suture in the south- best known (Fig. 2; Kilpeläinen, 1998). Most ern part of the VMB. of the volcanic and plutonic rocks were formed Nonetheless the arc-accretion hypothesis between 1.90 and 1.87 Ga (e.g. Huhma, 1986; does not readily explain the ubiquitous occur- Patchett & Arndt, 1986; Gaál & Gorbatschev, rence of plutonic rocks of similar age through- 1987). The volcanic sequences (including those out the Svecofennian Province, nor their exten- of the TampereSchist Belt (TSB) adjacent to sion far into the NW Karelian province. It has the VMB, Fig. 2) have been interpreted as hav- been recognised that the Svecofennian province ing island arc affinities, and models have been is unusual compared with Phanerozoic arc com- developed in which island arcs, formed over a plexes, and that the evolved nature of the volcan- series of subduction zones, have been progres- ics requires a pre-1.9 Ga history (Kähkönen et sively accreted to the Karelian craton (e.g. Park, al., 1989, p. 40; Nironen & Kähkönen, 1994). 1985; Ekdahl, 1993; Lahtinen, 1994; Nironen, Evidence for a pre-1.9 Ga component in both 1997). The province has come to be widely ac- the volcanic and plutonic 1.90–1.87 Ga suites cepted as an example of rapid crustal growth by has steadily accumulated (e.g. Lahtinen, 1994; arc accretion. Lahtinen & Huhma, 1997; Vaasjoki & Huhma, Twoprincipal episodes of accretion have 1999, for Finland; Valbracht et al., 1994; An- been proposed in Finland: at ~1.91–1.90 Ga, dersen & Sundvoll 1995; Claesson & Lundqvist east of the Central Finland Granitoid Complex, 1995; Billström & Weihed, 1996; Lundqvist et and at ~1.89–1.88 Ga, south of that complex al., 1998, for Sweden). (Lahtinen, 1994). The arc-accretionary model is It is also recognised that the pre-Svecofenni- reflectedin the terminology of the 1: 1 000 000 an ocean opened by 1.95 Ga (preferred by Niro- Bedrock map of Finland (Korsman et al., 1997): nen, 1997) and possibly by 2.1 Ga (Lahtinen, only the “Primitive arc complex of central Fin- 1994) or even earlier (Rutland et al., 2001b). land (1.93–1.87 Ga)”, immediately adjacent to Studies of detrital zircons in post-1.9 Ga meta- the Archaean of the Karelian Domain (C in Fig. sediments have revealed that source areas with 1; the Savo Belt in Fig. 2), is identified as old- ages of 2.1–1.9 Ga were available for erosion er than 1.90 Ga (see also Korsman et al., 1999). ~1.9 Ga ago (Huhma et al., 1991; Claesson et The VMB occupies a critical position between al., 1993; Welin et al., 1993), although no such the “Accretionary arc complex of southern Fin- areas could be identified.Lu ndqvist et al. (1998) land (1.90–1.82 Ga)” and the “Accretionary arc considered that, in Sweden, sediments were de- complex of central and western Finland (1.90– posited in the Bothnian Basin (Fig. 1) over an Pre-1.91 Ga deformation and metamorphism in the Palaeoproterozoic Vammala Migmatite Belt… 95 extended period from before 1.95 Ga to at least plexes were initiated as the fillof a large margin- 1.87 Ga. They recognised several phases of gra- al basin, when the contemporaneous active mar- nitic magmatism, but no accompanying defor- gin lay to the SW of the exposed Svecofennian mation or interruptions to sedimentation dur- Province (cf. Hietanen, 1975). Closure of that ing that period. Such evidence is not readily rec- basin was considered to be responsible for the onciled with an arc accretion process. principal episode of accretion and crustal thick- Lahtinen & Huhma (1997) argued that the ening. It was therefore proposed to restrict the isotopic and geochemical data for the Central term Bothnian Basin to the sequences deposit- Finland Granitoid Complex (CFGC) and TSB ed after the D 1 deformation and to apply Seder- (Fig. 2) indicated the presence of an evolved holm’s term ‘Svionian’ to the pre-D1 sequences. thick crust at 2.0 Ga. They proposed a tentative Dating the main period of deformation and model involving the presence of a continental metamorphism is the key to testing these two nucleus under the CFGC, surrounded by more models. If the Svecofennian was formed by arc juvenile island arc accretions. They inferred that accretion, then the main deformation and met- this older crust formed the basement to the old- amorphism should haveaccompanied or fol- est volcanic rocks (the Haveri Formation) in the lowed the deposition of the volcanic sequenc- TSB. Kähkönen (1999, p. 27) remarks that no es (e.g. at ~1.89–1.88 Ga for the VMB, Lahtin- preserved rocks of the proposed ~2.0 Ga crust en, 1994). If it formed by marginal basin accre- are known in central Finland. Nevertheless, the tion, the main deformation and metamorphism VMB remains a potential candidate for an older should havepreceded deposition of those se- basement (cf. Kähkönen, 1996). quences, i.e. before ~1.90–1.91 Ga. The present Early in workonthe Tampereareaitwas paper tests these alternative hypotheses for the suggested that an older Svionian Formation VMB in the classical Tampere area. was separated from a younger Bothnian Forma- tion by a great unconformity (Sederholm, 1897; 2.The Tampere Area 1931), i.e. the VMB formed basement to the TSB. This basement-cover hypothesis was aban- The Tampere area lies on the southern margin doned by later workers, who posited the strati- of the ‘Accretionary arc complex of central and graphic equivalence of the TSB and VMB (Kil- western Finland’, as shown on the 1: 1 000 000 peläinen, 1998, p. 7, and references therein). Bedrock map of Finland (Korsman et al., 1997; However, the Sederholm hypothesis has never see Figs.1 & 2). The Tampere Schist Belt (TSB) been convincingly refuted (Nironen, 1989, p. forms an asymmetric, upright, synclinal struc- 36; section 2.2 below), and the basement-cov- ture, about 20 km or less in width (Simonen, erhypothesis has recently been revived to ex- 1953; Nironen, 1989a), immediately south plain relations between the migmatite complex of the Central Finland Granitoid Complex of the Swedish Robertsfors Group and the Skel- (CFGC; see Figs. 2 & 3), and north of the Vam- lefte volcanic belt (Fig. 2). Rutland et al. (2001a mala Migmatite Belt (VMB). The main belt of and 2001b) argued that the early (D1 ) defor- volcanic rocks is less than 10 km wide and about mation and metamorphism in the Robertsfors 100 km long, while the TSB as a whole is about Group took place before deposition of the Skel- 160 km long. lefte sequence (i.e. before ~1.9 Ga). They also The predominant suite of weakly foliated suggested that pre-1.9 Ga complexes, forming granodiorites, tonalites and quartz diorites in basement to the volcanic sequences, are wide- the CFGC has been allocated on the map to the spread in the Svecofennian Province (see Fig. 2). 1.89–1.88 Ga interval, and similar rocks occur This led to an alternative plate tectonic model in in the TSB and VMB. Voluminous unfoliated which it was proposed that the basement com- granites in the CFGC, have been allocated to 96 Rutland,R.W.R.,Williams,I.S. and Korsman,K. Pre-1.91 Ga deformation and metamorphism in the Palaeoproterozoic Vammala Migmatite Belt… 97

Fig. 1. Regional settingand main elements of the Svecofennian orogenic province,and of the concealed Palaeopro- terozoic domains in the basement of the East European Craton (modified after Gee and Zeyen, 1996, and Claes- son et al., 2001). In the Svecofennian Province: Areas A, Band Care from Korsman et al., 1997: A,Accretionaryarc complex of southern Finland (1.90–1.87 Ga) B,Accretionaryarc complex of central and western Finland (1.90–1.87Ga); C, Primitivearc complex of central Finland (1.93–1.87 Ga).

Note that the Bothnian Basin has usuallybeen considered to extend over most of area B, which is nowlargelyoc- cupied by graniticrocks.TIB,Transscandinavian Igneous Belt (1.85–1.65Ga); Hi, Himanka; J, Jokkmokk; L, Luleå; R, Raahe; Sk, Skellefteå.

In the East European Craton,WLG, West Lithuanian GranuliteDomain; LEL, Latvian-East Lithuanian Belt; BBG, Be- larus-Baltic Granulite Belt; G, granulite belt, possible northwardextension of VG,Vitebsk Granulite Domain; CB, Central Belarus Belt; O, Okolovo Volcanic Belt (~2.0–1.98Ga); OMB, Osnitsk-Mikashevichi Igneous Belt (2.0–1.95 Ga); BG, Bragin Granulite Domain(~2.3–2.1 Ga);V-U,Volgo-Uralia ArchaeanProvince;TESZ,Trans-European Sutu- re Zone; M, Moscow.

Fig. 2. Geological mapshowing outcrops of Kalevian rocks on the SW margin of the Karelian Province and their proposed correlatives in the Svecofennian Province (modified after Rutlandetal., 2001b,Fig. 6). H, Himanka; M, Mikkeli; N, Nivala; Pih, Pihtipudas; P, Pyhajärvi; R, Robertsfors;T, Tampere; HF,Härnö Formation; RG, Robertsfors Group;TSB, TampereSchist Belt;VMB,Vammala migmatite Belt;VMC,Vaasa Migmatite Complex; WBB, Western Bothnian Belt; CFGC,Central Finland Granitoid Complex. Note that the pre-1.9 Ga metamorphic complexes, other than the Savo Belt, were previouslyconsidered to be partofthe Bothnian Basin but areheredistinguished as the SvionianBasin. 98 Rutland,R.W.R.,Williams,I.S. and Korsman,K.

~1.88 Ga. Some younger, more equidimensional bidites in the southern limb has yielded a SIMS granite plutons, dated at 1.88–1.87 Ga, occur age of 1898±8 Ma, and a high-K rhyolite near immediately north of, and within, the TSB (Fig. the top of this succession has given a TIMS age 3; see Nironen et al., 2000; Rämö et al., 2001). of 1892±3 Ma (Kähkönen et al., 2004). Both of these granite types occur across the The greywackes of the Tervakivi Formation, whole width of the CFGC (nearly 300 km) but above this rhyolite, terminate the previous up- they are not mapped within the VMB. ward-fining trend in the sediments. One con- South of the VMB (Fig. 15), the Hämeen- glomerate (the Veittijärviconglomerate) con- linna-SomeroVolcanic Belt (HSB; Hakkara- tains plutonic clasts dated at ~1.88–1.89 Ga inen, 1994) has similar characteristics to the (Nironen, 1989b), and provides evidence of TSB. Further south, the late Svecofennian gran- syn-depositional uplift and erosion. The depo- ite-migmatite zone (LSGM, Ehlers et al., 1993; sition age of this conglomerate must lie between Väisänen & Hölttä, 1999) is characterised by a the age of the underlying rhyolite (1892±3 Ma) distinctly younger (<1.84 Ga)suite of micro- and theage of overlying volcanics inthe core clinegranites and associated deformation and of the main syncline (1889±5 Ma, Kähkönen et metamorphism, superimposed on the 1.89– al., 1989). The latter date also further constrains 1.87 Ga early plutonic rocks and on the Kemiö- the age of the intrusion from which the grani- Mäntsälä volcanic belt (KMB). toid pebbles originated. Significant, and probably extensional, tectonic activity is also indicated by the change to basalts and andesites in the up- 2.1.The Tampere Schist Belt per volcanics; and by the substantial thickness of The detailed studies of the TSB outlined below arenites (Kalkku and Mauri arenites) and over- refer mainly to the main syncline on both sides lying mudrocks in the upper parts of the suc- of the lake north of Tampere (Fig. 3). Most of cession (Kähkönen, 1999, pp. 24–25). Accord- the ages quoted havebeen measured by ID- ing to Kähkönen (op. cit) the arenites were de- TIMS zircon U-Pb. The stratigraphic profilesin rived from a source similar to high-K rhyolites the syncline (Kähkönen et al., 1989; Kähkönen, at the top of the lower volcanics, implying that 1999) showanupwardtransition from tur- such volcanics were volumetrically more impor- bidite-dominated submarine environments to tant than now seen, and emphasising the impor- volcanic-dominated, in part subaerial, environ- tance of the syn-depositional erosion. ments. The turbidites below the volcanics occur Although the presence of conglomerates with mainly in the southern limb of the main syn- granitoid pebbles indicates tectonic activity dur- cline, where they are named the Myllyniemi for- ing deposition, the main compressional deformation (Fig. 3). The youngest detrital zircons mation of the TSB by folding appears to post- from two samples (Huhma et al., 199l; Claes- datethe whole stratigraphic succession (Niro- son et al., 1993) indicate maximum deposition nen, 1989a). This local F 1 , and associated struc- ages of “about 1.92–1.91 Ga” (Lahtinen et al., tures, are identified with the prefix ‘T’ below to 2002). distinguish them from deformation episodes in The lower volcanics, intercalated in the upper the VMB, which are identified with the prefix part of the turbidite succession, are mostly fel- ‘V’. The main TF1 folds are upright and some- sic. Kähkönen (1994) makes some analogy with what asymmetric with a near vertical axial plane

Stromboli shoshonitic lavas. Localities in the N schistosity,TS 1 .Noearlier schistosity is ob- limb of the main syncline givezircon ages of served, although the migmatites of the VMB to 1904±4 and 1898±4 Ma; ~1.89 Ga (Kähkönen the south do provide evidence of an older defor- et al., 1989). A unit considered to represent the mational-metamorphic event (op. cit. p. 34). lowermost major volcanic unit among the tur- The gross crosscutting relations of the plu- Pre-1.91 Ga deformation and metamorphism in the Palaeoproterozoic Vammala Migmatite Belt… 99

Fig. 3. Geological mapofpartofthe Tampereand Vammala belts (modified after Korsman et al., 1997), also showing keylocalities (see text).TSB,TampereSchist Belt; MM, Mauri Meta-arkose;VMB,Vammala Migmatite Belt; OF,Osara Formation; G1, granodiorite,tonaliteand quartz diorite (1.89–1.88 Ga); G2, granite (~1.88 Ga); G3, granite (1.88– 1.87 Ga). Note (1) the inferredpost-TSB shear zone in VMB, parallel to southern margin of TSB (cf. Fig. 5) (2) that the Haveri basalts and the overlying Osara formation areshown as partofthe TSB in accordance with published views.As discussed in the text, allocation to the VMB is herepreferred.

tons in the central TSB suggest that they were in the CFGC, has been dated at 1870±4 Ma emplaced late in the deformation history. How- (Nironen et al., 2000). ever,following detailed study of the structur- Prograde metamorphism, with the crystal- al relationships,Nironen (1989b) concluded lization of staurolite, garnet and andalusite, is that the plutons were syn-tectonic with TD1 . closely associated with TD1 . Andalusite contin- The measured ages of these plutons range be- uedtocrystallise after the deformation, when tween ~1.88 and 1.91 Ga, possibly because of sillimanite also began to crystallise, indicating the presence of some inherited zircon, and an that progressivemetamorphism continued af- age of ~1.88 Ga can be inferred for emplace- ter the deformation. Pressure has been estimat- mentand folding. The little-deformed Lavia ed at 3–4 kbar with peak temperatures of 470°C porphyritic granite (Fig. 3), a granitic body that in the central parts and 570°C in the margin- has been grouped with the post-kinematic suite al parts of the Tampere syncline (Kilpeläinen et 100 Rutland,R.W.R.,Williams,I.S. and Korsman,K. al., 1994). Thus, although the precise relations a“great discordance”between the TSBand between them are not clear, the metamorphism, VMB, as proposed by Sederholm, was wide- deformation and granitoid intrusion are closely ly discussed. But since the stratigraphic inves- related in time and constrained between ~1.89 tigations of Simonen (1953), “it has been large- and ~1.87 Ga. ly accepted that the areas arenot exotic relative to each other in either age or palaeogeog- raphy”.Kilpeläinen et al. (1994, p. 27) state, 2.2.The Vammala Migmatite Belt and relations “It was eventually concluded that the migma- to the TSB tite area was merely a marine part of the Tam- The belt of migmatised mica gneisses and pere island arc formation, in which the meta- schists south of Tampere (Figs. 2 & 3) trends E- morphic grade increases gradually from the syn- W and is up to 50 km wide. It is part of a belt clinally folded Tampere schist area to the anti- extending to east and west that has been given clinally folded Vammala migmatite area (Simo- various names in the literature, including Vam- nen, 1980).” Lahtinen (1996), on the basis of mala Migmatite Belt (Kähkönen et al., 1994), lithology and geochemistry, suggests broad cor- Pori-Vammala-Mikkeli Migmatite Zone (Kil- relations between the sediments of much of the peläinen, 1998), Mica gneiss-migmatite Belt VMB and those of the TSB. (Lahtinen, 1994; 1996), Psammitic migmatite The evidence for the stratigraphic correlation zone (Korsman et al., 1999), Tonalite Migma- between the VMB and TSB is not conclusive, tite Zone (Koistinen et al., 1996), and Tonal- however. Simonen (1953), in reviewing earli- ite-Trondhjemite Migmatite Belt (Mouri et al., er views critical of Sederholm’s hypothesis, did 1999). The name Pirkanmaa Belt has also re- not deny the existence of asub-Bothnian un- cently been recommended (Nironen et al., conformity but suggested (op. cit., p. 12) that 2002) for an extensive lithological-geographical it is the result of local erosion during orogen- area, the limits of which have not been fully de- ic evolution. This does not negate the possibili- fined.The belt is here considered to be one part ty that the metasediments of the VMB are old- of widespread pre-1.91 Ga metamorphic com- er than any part of the TSB. Nironen (1989, p. plexes in the Svecofennian domain (Fig.1). We 36) commented that Sederholm’s(1897) con- therefore revert to the geographically qualified clusion had not been satisfactorily refuted, and name, Vammala Migmatite Belt (VMB), to dis- noted that nothing was known about the age of tinguish the segment south of the TSB. the palaeosome in the migmatites. The possi- Geochemical studies (Lahtinen, 1996) show bility therefore remains that the VMB is strati- that the VMB has distinctivesediments with graphicallyolder than the TSB. However, the massive psammites, and some conglomerates, in basement-coverhypothesis also suggests that the NE, and pelites and black schists in the SW. the earliest deformation and metamorphism in Other greywackes in the VMB differ from those the VMB should pre-date the TSB. in the TSBinhaving agreater mafic compo- Aeromagnetic evidence, which was not avail- nent. Volcanic rocks occur mainly in the south- able to the early workers, provides supportto ern part of the VMB, and are mafic with both the basement-cover hypothesis. The volcanic se- N-MORB and E-MORB/WPB affinities, ap- quence in the main TSB syncline shows a sim- parently formed in a rift or marginal basin envi- ple linear pattern, consistent with the observed ronment (op. cit. p. 61). Some granitic and ma- TD1 folding. The greywackes immediately to fic intrusions also indicate a WPB component the south lack asignificant signaturebut are (op. cit. p. 91). known to sharethe same simple folding. The Kilpeläinen (1998, p. 7) noted that in the south-side-up shear zone, which separates the first half of the last centurythe possibility of TSB from the VMB (Fig. 3), is essentially a TD1 Pre-1.91 Ga deformation and metamorphism in the Palaeoproterozoic Vammala Migmatite Belt… 101

structure (Nironen, 1989a, p. 31) and there is ing D 3 as well as D 2 , and structures categorised an abrupt change to a strikingly more complex as D 2 may belong to two generations. aeromagnetic pattern in the VMB. This pattern TS1 is closely related to shear zones, and has been shown by Kilpeläinen (1998) to re- layering can be transposed parallel to TS1 flectmultiple folding. Upright folds of the earli- (Nironen, 1989a, pp. 27–31). In the VMB, the est schistosity have variable trends as a result of corresponding deformation appears to be con- refolding on axes approximately parallel to the centrated in broad major shear zones trend- main folds of the TSB (Fig. 5). ing ENE in the area south of Tampere, WNW Kilpeläinen (1998) has provided a compre- through the Koosanmaa sub-area, and NNW hensivediscussion of the structures in various through the Vammala area (compare Figs. 3 & sub-areas of the VMB (see Fig. 5). He notes the 5 with Fig. 4). Within these zones, and especial- difficultiesin making regional correlations when ly in the ENE zone, the granitoid intrusions are the intensity of deformation episodes varies and elongated parallel to the zones and to the post- when the relative and absolute ages of peak met- TSB structures (e.g. in the Stormi sub-area of amorphism also vary. However, he allocated all Fig. 5, detailed in Kilpeläinen, 1998, Fig. 16, the structures to three generations distinguished and in the SE corner of the Pirkkala sub-area). on the basis of style and superposition (op. cit., Folds of VF2 style may also be near parallel to pp. 74–81): the zones, perhaps as a result of rotation. More-

over, transposition of layering parallel to VS2 or

–D 1 , represented throughout by a penetrative VS3 may have resulted locally in such schistosity

layer-parallel biotite schistosity,inferred to being indistinguishable from VS1 . Outside the have been near horizontal when formed. major shear zones, the granitoid intrusions are

–D 2 ,represented by upright, near isoclinal, of irregular form and their contacts transect the

macroscopic folds, responsible for the present trends of folds of VF2 style. The post-TSB de-

near vertical attitude of S 0 /S1 . S 2 may occur as formation is represented by aVS 3 crenulation an axial plane schistosity or as a crenulation cleavage (e.g. in the Myllymaa, Ellivuori and

of S 1 . Hämeenkyrö sub-areas of Fig. 5, detailed in Kil-

–D 3 ,represented by open upright folds and peläinen, 1998, Figs. 8, 12 & 31). shear zones. They fall into two, probably Peak metamorphism was considered to have

conjugate sets, trending ENE to NE, and occurred during VD1 (op.cit., pp.96–97), WNW to NW (Fig. 5). with sillimanite, garnet and cordierite growth in appropriate lithologies. However, crystallisa- As noted above, we now prefix‘ V’ to these cate- tion of garnet always preceded that of cordier- gories. The two VD3 shear directions correspond ite, and cordierite often clearly post-dates the well in trend with the TS1 schistosity, axial plane VD1 deformation,probablyreflecting decom- to the main post-TSB folds, and our examina- pression. Metamorphism and migmatisation tion of critical fieldevidence below suggests that in the VMB has recently been dated, using sev- their correlation is generally valid. This is in ac- eralisotopic methods, at ~1.88 Ga (Mouri et cord with the interpretation of Koistinen et al. al., 1999). It was inferred (op. cit.) that these (1996, p. 49), although Kilpeläinen (1998, Fig. migmatites were metamorphosed at peak con-

49, p. 75, pp. 78–79) correlates VD2 with the ditions of 700–750ºC and 4–5 kbar, and also main (TD 1 ) folding in the TSB, as shown in Fig. recordaretrogressivephase involving decom- 5. We suggest that problems of correlation have pression and cooling to around 500–650°C and arisen because the above scheme, while general- 3–4 kbar, in the andalusite stability field.We ar- ly applicable, is somewhat oversimplifiedin that gue below (section 7.4) in the light of our data, new planar schistosity may develop locally dur- that the metamorphism dated by Mouri et al. is 102 Rutland,R.W.R.,Williams,I.S. and Korsman,K.

Fig. 4. Aeromagnetic low-altitude map of the Tamperree Vammala area. Geological Survey of Finland, Espooo.. Compiled by Maija Kurimo (after Kilpeläinen, 1998, Fig.ig. 3, p.17). Same area as Fig. 5.

the post-TSB (TD 1 ) metamorphism and not the calityV-4 in Fig. 3). Thus in both the TSB and earlier VD1 . VMB, episodes of deformation and metamor- A date of 1881±4 Ma for a quartz diorite in phism synchronous with granitoid intrusion the central VMB has been interpreted as the at ~1.88 Ga have been identified. In the VMB age of magmatism synchronous with the local there is an earlier episode, VD1 , thought to coin-

VD2 folding (Kilpeläinen, 1998, pp. 41–42; lo- cide with peak metamorphism, but no such epi- Pre-1.91 Ga deformation and metamorphism in the Palaeoproterozoic Vammala Migmatite Belt… 103

Fig. 5. MetaMetamorphicmorphic mapop offtthehe Tampere-Vammala area (from Kilpeläinen, 1998, Fig. 51, p. 79). Mapsp sheets:heets: 2122 = Ikaalinen, 2124 =Viljakkala-Teisko, 2121 =Vammala, 2112 =Huittinen, 2114 =Toijala. sode has been identifiedin the TSB. Kilpeläinen peläinen et al. (1994, p. 30) explain this only by et al. (1994, p. 30) also recognised that the ear- suggesting that “the temperatureand age of met- ly, originally recumbent (VD 1 ) deformation and amorphism and the beginning of structural evo- associated garnet-cordierite-sillimanite-biotite lution”increase as afunction of depth. This and metamorphism are characteristic of the VMB, other tectonic interpretations depend heavily on while the TSB was metamorphosed only during the assumption that the stratigraphic correlation the later upright (TD 1 ) deformation and is char- between the turbidites of the TSBand the meta- acterised by andalusite, staurolite and garnet as- sediments in the north of the VMB is valid (e.g. semblages (Fig. 5). As a result, in the VMB, VS2 Kilpeläinen et al. 1994, pp.30–32; Kilpeläinen, and VS3 areusually crenulation cleavages, su- 1998, pp.77, 99, 101 and 113). On this assump- perposed on VS1 , while in the TSB, TS1 forms tion, VD1 must be younger than at least partof a penetrative schistosity, with no evidence of an the TSBand therebywas assumed to haveamax- earlier schistosity. imum age of ~1.91 Ga,the age indicated by the Thereisnodoubt from the field evidence that youngest detrital zircon in the TSBturbidites,

VD1 is earlier than the intrusivegranitoids, but it and aminimum age of ~1885 Ma,the age of in- has previously been considered to be close in age trusions associated with the main TSBfolding. to the main TSBfolding, and to post-date dep- An alternative reading of the evidence would osition of all or partofthe TSBsuccession. Kil- argue that granitoid pebbles dated at ~1.89 Ga 104 Rutland,R.W.R.,Williams,I.S. and Korsman,K.

were incorporated in the TSB succession before ty that we attribute to VD1 .Inthe north, V F 3

1889±5 Ma. Given the field evidence that VD1 crenulation cleavage and open folds trending pre-dates such granitoids, VD1 must haveoc- WNW,are parallel to the trend of the syncline curred before the deposition of the upper part of TSBrocks to the northeast and the TSBsyn- of the TSB succession; and given the fact that cline north of the Lavia Granite. This suggests a the lower part of the TSB succession shows only correlation between this cleavage and the princi- the same post-TSB structures as the upper part, pal folding (TD 1 )inthe overlying TSB. it is also plausible that VD1 pre-dates the whole In the extreme SE of the map sheet, this of the TSB succession. If this is correct, the pre- gneiss complex is juxtaposed with another syn- viously assumed stratigraphic correlation of the cline of TSBrocks, trending NE. These are sedimentation in the VMB with that in the TSB muscovite chlorite grade metasediments (Fig. is not possible and no longer places an older 5), which preserve sedimentarystructures and limit on VD1 . younging directions (Ikaalinen map sheet, 1993 In general, early structures, which might cor- edition, locality I-2 in Fig. 3), in remarkable relate with VD1 , are absent from the TSB, con- contrast to the nearby high-grade gneisses that sistent with VD1 being older, and with the base- are much coarser grained and show evidence of ment-coverhypothesis. However, some out- superposed deformations. Other gneisses in the crops containing such structures have previous- vicinity have more variable E to NE trends and ly been allocated to the TSB (Koistinen et al., display adifferentiation schistosity,and near

1996; Kilpeläinen, 1998). We have therefore re- isoclinal VF2 folds plunging steeply southwest. examined the critical fieldevidence, especially in Open VF3 folds and crenulation of the early fo- key outcrops close to the boundary between the liation and lineation have a NE to ENE trend TSB and VMB. In no case does it appear that and can be correlated with the trend of the ad- their allocation to the TSB is valid. jacent TSBsyncline. We interpret the abrupt change in structural and metamorphic character (compare Fig. 32b with Figs. 32c & 32e of Kil- 2.3. Critical field evidence peläinen, 1998) as indicating a relatively undis- The main gneiss area in the SE of the Ikaalin- turbed unconformable relationship. en map sheet (Virransalo &Vaarma, 1993; the A key area lies on the southern margin of the Hämeenkyrösub-area of Kilpeläinen, 1998; see TSB, immediately south of the outcrop of the Figs. 4&5)has sometimes been interpreted as Mauri meta-arenite (Fig. 3). Kilpeläinen (1998, partofthe TSB, although its broad similarity to p. 99) states that D 1 structures are visible in the the Vammala migmatites has usually been recog- andalusite schists above the Mauri meta-arenite, nised (for example, Kähkönen, 1999; Huhma et which itself is high in the stratigraphic succes- al., 1952). It appears to form adome-like struc- sion (section 2.1). These schists were inferred to ture, with synclinal prolongations from the main overlie “the western contact of the north-south TSB outcrop both to the NE and SE. The aero- trending partofthe meta-arkose”.Atlocality magnetic pattern and the fold pattern showsim- V-1(Fig. 3; GR 6819.80N, 2460.30E, Vam- ilar complexity to that of the main partofthe mala 1:100 000 map sheet, Matisto, 1967), VMB (Kilpeläinen, 1998, pp.57–62, and Fig. the schists show growth of andalusite after the 31), suggesting that this is anortherly extension principal fine-grained schistosity (op.cit. Fig. of the VMB. Oneofour samples, Ikaalinen-1 (I- 7b), which can be correlated with TS 1 . No ear- 1inFig. 3) was thereforetaken from this gneiss lier schistosity or fold structure has been identi- area. VF2 folds in potassium feldspar-sillimanite fiedin this vicinity, and there can be little doubt gneisses haveasinuous, roughly NE, trend (Fig. that at this locality the schists do stratigraphi- 5) and refold an early layer parallel schistosi- cally overlie the Mauri meta-arenite. However, Pre-1.91 Ga deformation and metamorphism in the Palaeoproterozoic Vammala Migmatite Belt… 105 rocks several kilometres to the SSW that have ing of the Mauri meta-arenite. The sharp con- also been allocated to the TSB show quite dif- tact with the meta-arenite is hereinterpreted ferent structural relationships. as a TS1 shear parallel to the TS1 schistosity in

At locality V-2(Fig. 3), refolding of an earlier the latter, and to the correlated VS3 crenulation

(local S 1 )schistosity has been recognised (op.cit. cleavage in the Vammala migmatites. The aer- Figs. 4, 5a and 5b; GR 6817.70N, 2453.65E). omagnetic map (SW corner of sub-area 1, Fig. These rocks were included in the TSBonthe as- 4) defines the location of the W trending shear, sumption that they also stratigraphically overlie which separates the contrasting aeromagnet- the Mauri meta-arenite. Ourdetailed re-exami- ic signatures of the Vammala migmates to the nation shows, however, that this is highly unlikely south and the Mauri meta-arenite to the north. and that this locality is separated from the meta- The corresponding contrast between muscovite/ arenite by an important shear zone. sillimanite/potassium feldspar and andalusite/ The meta-arenite shows good primary layer- staurolite metamorphism is shown in Fig. 5 (af- ing, which is often planar.Locally,tight upright ter Kilpeläinen, 1998, Fig. 51). mesoscopic folds occur with afine-grained axial- The Mauri arenites are considered to be high plane schistosity that has constant orientation. As in the stratigraphic succession. Their presence at locality V-1, this single upright schistosity cor- on the southern margin of the TSB, and in di- relates with the main folding of the TSB(i.e. it rect, albeit sheared, contact with the migmatites isTS 1 ). The well-preservedplanar layering and of the VMB, strongly suggests erosion and over- the simple prograde metamorphism preclude the step of the lowersuccession as aresult of the possibility that earlier structures were present, but syn-depositional tectonic activity. Thus, prior to obliterated by transposition. the folding and shearing of the TSB succession, The refold structures at locality V-2 contrast it is likely that the Mauri arenites rested uncon- sharply with this simple structure in the adja- formably on the already deformed VMB. cent Mauri meta-arenite. The rocks display rela- Thereisanother key area west of Pirkkala tively coarse S 1 and S 2 schistosities that were de- on the Tampere map sheet. West of locality Ta- veloped during high-grade metamorphism pri- 1(Fig. 3), greywackes with good sedimentary or to refolding by F3 (Kilpeläinen, 1998, Figs. structures and N-S strike exhibit a simple aero- 5a and 5b in strong contrast to Fig. 7b above). magnetic signature showing distinct N-trending Locally, these coarse–grained rocks contain sil- bands (Fig. 4), in strong contrast to the Vam- limanite (op. cit. p. 20). Granitic veins cut the mala migmatites to the east, which show refold- early schistosity but behaved as competent lay- ing on a NW trend. Again the structural change ers during the refolding. F 3 is accompanied by is accompanied by an abrupt change in meta- an axial-plane crenulation cleavage that is paral- morphic assemblages from andalusite + stauro- lel to the single prograde schistosity in the Mau- lite to sillimanite + potassium feldspar + mus- ri meta-arenite. This youngest cleavage is very covite (sub-area 7 in Fig. 5). Locality Ta-1 (Fig. strong close to the meta-arenite. 36a of Kilpeläinen, 1998) is here interpreted as We therefore now correlate the single schis- a shear zone between fine-grainedTS B rocks to tosity (TS 1 )inthe meta-arenite with the F 3 the west (op. cit. Fig. 36c) and coarse-grained crenulation cleavage in the V-2 outcrops and re- retrograded VMB gneisses to the east (op. cit. gard the latter as part of the Vammala migma- Fig. 38). tites. We now consider the S 1 and S 2 schistosities These examples from critical localities con- in the V-2 outcrops to be VD1 and VD2 struc- sistently demonstrate the very abrupt change tures that havebeen re-folded and retrograd- from coarse-grained gneisses, showing com- ed, producing abundant muscovite, during the plex folding and an early VD1 metamorphism,

VD3 deformation that produced the TF1 fold- to much finergrained schists showing only sim- 106 Rutland,R.W.R.,Williams,I.S. and Korsman,K. ple post-TSB structures and a late lower grade cession no. 99 0573; ) was taken from the inte- metamorphism associated with the TD1 schis- rior of the migmatite belt, in the NE of the Toi- tosity. The latter consistently correlates with a jala 1:100 000 geological map area (sheet 2114, late fold and crenulation cleavage phase, VD3 , Matisto, 1973). Horizons interpreted as black associated with retrogression in the gneisses. schists are prominent in the aeromagnetic im- Thus, the peak metamorphisms on either side age of the area and appear to delineate an ear- of the boundaries at these localities are geolog- ly macroscopic fold structure (local F 1 ), refolded ically of different age, as has long been recog- (by the local F 2 ) on a NE trend (Fig. 4). The ge- nised (Kilpeläinen et al., 1994, p. 30, see above). ometry suggests an original NNW trend for F 1 . Only the layering in the TSB rocks is parallel At the sample locality (Fig. 3; GR 6796.35N, to the boundaries. The boundaries arethere- 2492.65E; 7.3km from Lempäälä towards Kär- fore discordant with respect to lithology, meta- jenniemi on road 3041), in the NW limb of this morphism, and structural style and history. The major structure, the principal schistosity (S1 ) is contacts must be abrupt and the relationships sensibly parallel to primary lithological banding are quite characteristic of a basement-cover re- (Fig. 6a), and is allocated to VD1 . The sample lationship in which the basement shows defor- is a plagioclase-quartz-biotite gneiss dominated mation and metamorphism not represented in by fine-grained quartz and plagioclase (~0.1–0.2 the cover. The presence of the antiformal inlier mm), with biotite and minor garnet. There are of VMB in the SE Ikaalinen area, with synclinal also larger (~1 mm) plagioclase grains with nu- prolongations of the TSB to north and south, merous quartz inclusions, as well as large quartz (andreflected in the aeromagnetic contrast in grains and aggregates, and elongate biotite ag- Fig. 4) also strongly supports the basement-cov- gregates, both up to 2 mm long. er hypothesis. The post-TSB metamorphism has The metamorphism in this part of the VMB a maximum age of 1889±5 Ma, the age of the is characterised by sillimanite +garnet (Kil- youngest dated volcanic member, and probably peläinen et al., 1994). Porphyroblast-schistosi- peaked at ~1880 Ma. We have therefore tried ty relations suggest that peak metamorphism oc- to date the VD1 metamorphism directly, to pro- curred during VD1 (Kilpeläinen, 1998, Fig. 61), vide a definitive test of the hypothesis. although metamorphic crystallisation and migmatisation continued in some areas throughout 3. Sample selection the structural evolution (e.g. Kilpeläinen, 1998, pp. 55–56). At the sample locality, thin bands

Korsman and Korja (1994) remarked, “Un- of concordant leucosome are parallel to S 1 and fortunately, we have very little information on probably formed during VD1 .Younger leuco- the evolution of the orogeny in the time in- some occurs in crosscutting veins, which lo- terval 1980–1885 Ma because the major ther- cally form a conjugate pair, suggesting that the malpulse at 1885 Ma obliterated and melted younger generation was emplaced under more the rocks recording the pre-collision orogenic brittle conditions. Neither leucosome genera- stages.” Therefore, for the present study an at- tion was visible in the collected sample. tempt was made to select samples containing Granitoid sample Toijala-2 (ANU accession

VS1 where the visible effect of younger events no. 99 0575) was collected about 1.1km NW of was minimal, and where the S 2 component in Toijala-1, from a major road cut west of Mäy- the schistosity was likely to be minimised. häjärvi, about 6.1km from Lempäälätowards Kärjenniemi (GR 6797.15N, 2491.80E). The rock is quartz-dioritic gneiss without micro- 3.1.The Toijala area cline, but is cut by undeformed leucocratic mi- Metasedimentarysample Toijala-1 (ANU ac- crocline bearing veins. Large rafts of the coun- Pre-1.91 Ga deformation and metamorphism in the Palaeoproterozoic Vammala Migmatite Belt… 107

Fig.6.Field photographs to illus- trate structural relationships. a. GreywackeatToijala-1 sample locality.Note that early leucosomeand schistosity areparalleltoprimarygar- netiferous layer(above pen- knife). b. Enclave of metasediment in granitoid at To ijala-2 sample locality, showing that emplacement post-dated the earlyschistosity and leucosome. tryrockoccur in thegranitoid and it is clear en 1:100 000 geological map area (sheet 2122, that the main schistosity and concordant leuco- Virransalo & Vaarma, 1993), about 8km W of some in the rafts pre-date the emplacement of Hämeenkyrö (Fig. 3). The locality was previous- the granitoid (Fig. 6b). The macroscopic field ly studied by Kilpeläinen (1998, pp. 58–59 and relations also suggest that granitoid emplace- Fig.32a, GR 6837.00N, 2449.50E). ment post-dated the macroscopic refolding (lo- The general structurehas been outlined in cal F 2 ) in the country rock, referred to above. section 2.3 above. The area is characterised by Thus the age of granitoid emplacement provides the potassium feldspar +sillimanite ±musco- a minimum age for VF1 and probably for the lo- vite metamorphic assemblage (Fig. 5 after Kil- cal VF2 also. peläinen, 1998, Fig. 51). The sillimanite oc-

curs as fibrolite aggregates, elongated in S 1 , and commonly shows retrogression to musco- 3.2.The SE Ikaalinen area vite (op. cit. p. 61). We infer that peak meta-

Metasedimentarysample Ikaalinen-1 (ANU morphism occurred during D 1 and that the ret- accession no. 99 0574) comes from the Ikaalin- rogression probably occurred during the post- 108 Rutland,R.W.R.,Williams,I.S. and Korsman,K.

TSB events, when the rocks of this area were at (1996). A primary ion beam of ~3 nA, 10 kV no great depth below the overlying TSB. Ikaa- O2– was focused to a probe ~20 µm diameter linen-1 is acoarse-grained arkosic metasedi- and the sputtered secondaryions extracted at ment, resembling adeformed granitoid. The 10 kV, mass analysed at a resolution of ~5000, larger grains are dominated by quartz, but pla- and the Ce, Zr, Pb, U and Th species of interest gioclase, microcline and fine-grained rock frag- measured on a single electron multiplier by cy- ments rich in apatite also occur. A strong schis- clic stepping of the analyser magnet field.We ak tosity (S1 ) is approximately parallel to the pri- energy filtering was used for the monazite anal- mary layering, dipping gently SSW, and a line- yses to reduce a molecular isobaric interference ar structure trends NNW. F 2 folding is not seen at mass 204. Isotopic compositions were meas- at this locality, but is prominent elsewhere (op. ured directly, without correction for mass frac- cit. pp. 58–59). tionation. Common Pb levels were comparable to those measured on Pb-free targets, so were 204 4. Analytical techniques corrected, using Pb, assuming laboratory contamination by Pb of Broken Hill galena compo- About 500 g of each rock sample were crushed sition. Corrections for Pb-U fractionation uti- to <250 µm and the high-density minerals >20 lised apower law (Claoué-Long et al., 1995). µm diameter extracted using density and mag- The correction factor for Th-U fractionation in netic separation procedures designed to prevent monazite (0.863) was calculated directly from a inter-sample cross contamination and sampling 208Pb/206Pb-Th/U isochron as outlined by Wil- bias. Final purification was by hand picking. liams et al. (1996). Analytical uncertainties list- Zircon and monazite were mounted separately ed in the tables and plotted in the figures are in epoxy resin with grains of reference materi- one standard error precision estimates based on al; zircon AS3 (radiogenic 206Pb/238U = 0.1859) the uncertainties in the time-interpolated count and SL13 (238 ppm U), and monazite Thomp- rates from which the ion ratios were calculated. son Mine WB.T.329 (radiogenic 206 Pb/238 U= Ages were calculated using the constants recom- 0.3152, 2100 ppm U). Mounts were polished mended by the IUGS Subcommission on Geo- to expose the crystal interiors, photographed in chronology (Steiger & Jäger, 1977). Uncertain- transmitted and reflectedlight, then washed and ties in the ages cited in the text are 95% confi- Au coated before analysis. Prior to analysis, the dence limits, namely t σ , where t is ‘Student’s t’. zircons also were imaged by SEM cathodolumi- nescence (CL) to document their internal zon- 5. Zircon isotopic compositions ing. These images were obtained using an Hi- tachi S-2250N SEM fittedwith an internal par- 5.1. Quartz diorite Toijala-2 abolic mirrortoboost signalstrength. Mona- The zircon from this gneissic granitoid consists zite growth structures were imaged by backscat- of mostly euhedral, finelyzo ned prismatic grains tered electron (BSE) after analysis, using a Cam- 50 to 150 µm diameter with aspect ratios of 2: bridge Instruments Stereoscan 360 fitted with a 1 to 5:1, features typical of zircon from a fel- 4-quadrant solid state BSE detector, and operat- sic plutonic rock. The CL of the zircon grains is ed at 20 kV with a 10 nA probe current. Imag- very weak, but shows that, although most grains es from both SEMs were recorded digitally, and are euhedrally zoned throughout, some contain the brightness and contrast subsequently opti- a structurally discordant core, and some have a mised off-line using Adobe Photoshop™. thin, discordant, more luminescent rim. U-Th-Pb was analysed on the RSES Fourteen areas were analysed for U-Th-Pb, SHRIMP II ion microprobe using procedures including two on cores and one on a CL-bright similar to those described by Williams et al. rim (Table 1). The zircon has a very wide range Pre-1.91 Ga deformation and metamorphism in the Palaeoproterozoic Vammala Migmatite Belt… 109

Fig. 7.Te ra-Wasserburg concordia diagram forzircon from the To ijala-2 quartz-diorite. of U contents (70–2070 ppm), the lowest value es and morphologies. Most grains arestubby, being from one very strongly luminescent core subhedral prisms 70–100 µm diameter,but (12.1). With the exception of that core, Th/U is some more equant grains reach 200 µm. Tr uly very uniform (mostly 0.2–0.3), consistent with euhedral grains are rare. Some cores and euhe- all the grains sharing a common igneous origin. dral growth zoning arevisible in transmitted Again with the exception of that core, all the light, but most grains appear relatively struc- analyses plot in a uniformly discordant cluster tureless and inclusion-free. Microfracturing is at ~1.9 Ga (Fig. 7). The analyses of the rim and common, however, many grains being cut by the other core plot within this group. There is a a maze of fine cracks. Such grains were not an- significant dispersion in radiogenic 206Pb/238U, alysed. The CL of the zircon is stronger on av- so the weighted mean of the 207Pb/206Pb meas- erage than that from the quartz diorite and var- urements, 0.11555±0.00014 ( σ ), all of which ies more in strength both within and between areequal within analytical uncertainty,gives grains (Fig. 8). About 60% of the grains have the best estimate of the zircon age, 1888±5 (tσ ) very weak CL. These consist predominantly of Ma. This is interpreted as the age of crystallisa- finely euhedrally-zoned zircon, sometimes with tion of the quartz diorite magma. It is compara- discordant cores and/or overgrowths. The re- ble to the age of granitoid clasts that appear in maining grains have much stronger CL, and fall conglomerates below the upper part of the suc- into two main groups, those with fine euhedral cession in the TSB, by which time such grani- zoning, and those in which the zoning is broad- toids were exposed to erosion. The luminescent er and less regular. All these commonly have dis- core is significantlyolder , 1992±23 ( σ ) Ma, and cordant cores and overgrowths. There are two is interpreted as being inherited from the quartz types of overgrowths, CL bright and CL dark. diorite’ssource or host rock. Lahtinen (1996, Both occur as thin structureless layers seldom pp.92–93) notes only one U-Pb zircon age >20 µm thick, mostly on the CL bright grains from the VMB, a granodiorite from Valkeako- and in about equal abundance. Where both oc- ski giving 1896±16 Ma, indistinguishable from cur on the same grain, the dark overgrowth is the age of Toijala-2. usually outermost. Fifty-two U-Th-Pb analyses were made of 39 grains (Table 1). Most attention was focused on 5.2. Paragneiss Toijala-1 the overgrowths to try to date post-sedimenta- Zircon from Toijala-1 has a wide range of siz- tion metamorphism. Some host grains to the 110 Rutland,R.W.R.,Williams,I.S. and Korsman,K.

Fig. 8. Catholuminescence imagesofanalysed zircon grains from the Vammala Metamorphic Belt. – Rows 1and 4: LowTh/U overgrowths formed during the ~1.92 Ga metamorphism from To ijala-1 (grains31, 34, 28) and Ikaalinen-1 (grains 1, 9), and associated slightlyolder anomalous grains(grains T- 1, 13, 28 and I-1, 14, 19; see text).Host grains (e.g. I -1, 31, 34 and T- 1, 1, 9provide examples of the main populations of zoned grains. – Rows 2and 5: overgrowths and associated grains representativeofthe ~1.98 Ga metamorphism from To ijala-1 and Ikaalinen-1 (grainsT-1, 27, 35, 23, 20, 8and I-1, 2, 29, 4, 12, 15; see text). – Row3:examples of Archaean grains with overgrowths from Ikaalinen-1. Pre-1.91 Ga deformation and metamorphism in the Palaeoproterozoic Vammala Migmatite Belt… 111

Fig. 9.Th (ppm) vs U(ppm) forthe To - ijala-1 paragneissand the Ikaalinen-1 arkosic gneiss as afunction of crystal zoning patterns. overgrowths were also analysed, as well as select- The isotopic compositions fall into three ed examples of the main grain types. The zir- main groups, a small number of grains at ~2.7 cons have a very wide range of U contents (17– Ga,and two major populations at ~1.98 and 1860 ppm) and Th/U (0.007–1.8), consistent 1.92 Ga (Fig. 10). Most analyses are near con- with zircon formed under a variety of geological cordant, but in each population there is signifi- conditions. The main population of euhedral- cant dispersion in Pb/U, with a limited range of ly-zoned zircon (Fig. 9) has on average the low- radiogenic 207Pb/206Pb. est U (40–760 ppm) and highest Th/U (>0.4). The Archaean group is comprised of analyses U tends to be higher (>200 ppm) in both the of one zoned grain and three overgrowths. The structureless zircon and overgrowths. A few of latter were analysed in search of younger meta- these grains have moderate Th/U similar to the morphic overgrowths, from which they could zoned grains, but in the majority,dominant- not be distinguished by either appearance or ly overgrowths, Th/U is much lower. Many of CL response. Twoare CL bright and one CL these have Th/U below 0.1, a feature common- dark. None of the analysed areas is low Th/U. ly found in zircon grown in metasediments un- The four analyses also have the same radiogen- der high-grade conditions (Williams & Claes- ic 207Pb/206Pb within analytical uncertainty, giv- son, 1986; Heaman et al., 1990; Williams et al, ing a weighted mean age of 2715±15 (tσ ) Ma. 1996). These grains indicate that Archaean rocks, some 112 Rutland,R.W.R.,Williams,I.S. and Korsman,K. 6 ± 4 7 9 9 9 8 6 5 6 9 7 6 8 8 11 11 11 14 11 11 12 23 11 35 11 57 17 1885 1890 1889 1892 1885 1887 1903 1874 1895 1886 1888 1881 1905 2710 1992 2719 2703 2060 2737 1964 1927 1938 2016 1982 1988 1978 2001 1914 07/206 a) ±2 12 12 14 17 16 15 17 17 19 23 18 15 14 25 37 17 31 17 12 13 15 14 37 11 20 16 18 101 ent Ages (M ent Ages (M 1857 1872 1828 1847 1830 1836 1874 1849 1859 1848 1852 1857 1814 2684 1868 2503 2750 2040 2701 1821 1880 1848 1934 1705 1954 1788 1777 1931 06/238 par ppar Ap ±2 th. 19 15 23 27 34 32 34 29 43 44 40 26 47 38 56 25 46 27 21 20 22 37 34 23 31 152 100 272 ow rgr ve 1884 1867 1830 1855 1841 1849 1877 1823 1869 1818 1871 1935 1763 2693 1962 2479 2743 2064 2763 1868 1871 2138 1929 2258 1975 1921 1850 1942 208/232 Th/U o w 0.3 0.2 0.4 0.5 0.5 0.6 0.6 0.5 0.6 0.8 0.6 0.4 0.6 0.7 1.3 0.3 0.7 0.3 2.1 0.3 0.6 3.1 0.5 0.4 0.4 1.0 0.5 0.5 ±% O - Lo Pb Pb*/ 206 207 0.11530 0.11564 0.11560 0.11581 0.11532 0.11546 0.11649 0.11464 0.11600 0.11537 0.11549 0.11505 0.11662 0.18628 0.12244 0.18738 0.18550 0.12725 0.18937 0.12052 0.11802 0.11875 0.12412 0.12175 0.12213 0.12150 0.12306 0.11723 % 0.8 0.7 0.9 1.1 1.0 0.9 1.0 1.1 1.1 1.4 1.1 1.0 0.9 1.1 2.3 0.8 1.4 1.0 4.5 0.7 0.8 1.0 0.9 2.5 0.7 1.3 1.0 1.1 eless rim; L uctur U± Pb*/ tr 238 206 0.3338 0.3369 0.3278 0.3317 0.3282 0.3295 0.3374 0.3322 0.3343 0.3319 0.3327 0.3339 0.3250 0.5163 0.3361 0.4744 0.5321 0.3723 0.5204 0.3264 0.3386 0.3320 0.3498 0.3027 0.3540 0.3197 0.3175 0.3493 1.0 0.8 1.3 1.6 2.0 1.8 2.0 1.7 2.4 2.6 2.3 1.4 2.9 1.5 3.0 1.1 1.8 1.4 5.9 1.1 1.1 4.9 1.2 1.9 1.8 1.3 1.7 ±% e; SR - S 12.7 Finland Th Pb*/ eless cor 232 208 0.0977 0.0968 0.0947 0.0961 0.0954 0.0958 0.0973 0.0944 0.0969 0.0941 0.0970 0.1005 0.0911 0.1425 0.1019 0.1305 0.1454 0.1075 0.1465 0.0969 0.0970 0.1116 0.1002 0.1182 0.1027 0.0997 0.0958 0.1009 uctur southern tr 0.6 0.3 0.8 0.9 1.4 1.5 1.4 1.2 1.9 1.9 1.8 0.8 2.6 0.8 1.3 0.6 0.8 0.9 2.3 0.8 0.6 4.3 0.7 6.8 1.2 1.2 0.6 1.1 ±% . belt, e; SC - S Pb Pb*/ 206 208 0.06619 0.19581 0.07493 0.06444 0.05744 0.09811 0.09368 0.08476 0.06220 0.06132 0.06625 0.07866 0.03915 0.50265 0.28205 0.21053 0.33747 0.13700 0.32150 0.07535 0.30174 0.19176 0.24222 0.23946 0.22301 0.15729 0.24416 0.14717 migmatite oned cor Pb 206 mmala C - Z 0.003 0.012 0.069 0.014 0.025 0.134 0.034 0.094 0.106 0.068 0.036 0.044 0.211 0.053 0.016 0.016 0.083 0.080 0.536 0.039 0.011 1.414 0.011 0.090 0.100 0.296 0.032 0.011 % comm.# Va th; Z the ow Pb Pb/ face common Pb using 204Pb om 204 206 ergr fr 5.0E+05 1.4E+05 2.3E+04 1.3E+05 6.7E+04 1.2E+04 4.8E+04 1.7E+04 1.5E+04 2.4E+04 4.5E+04 3.6E+04 7.6E+03 3.0E+04 1.0E+05 1.0E+05 1.9E+04 2.0E+04 3.0E+03 4.2E+04 1.4E+05 1.1E+03 1.4E+05 1.8E+04 1.6E+04 5.4E+03 5.0E+04 1.4E+05 ed sur con U .24 .14 Th/ zir 0.23 0.68 0.26 0.22 0.20 0.34 0.32 0.30 0.21 0.22 0.23 0.26 0.14 1.82 0.93 0.77 0.47 0.25 1.05 0.57 0.85 0.61 0.77 0.50 0.81 0.51 of y-deriv eless; O - Ov 91 01 95 86 59 48 50 53 66 Th 469 209 109 112 106 144 166 904 360 172 569 298 368 228 261 159 236 117 ppm 1079 yses uctur . tr anal 66 72 U 91 71 806 489 479 331 328 288 277 223 221 549 380 758 676 540 523 435 372 340 316 292 230 ppm 2073 1583 1182 oijala 2) oijala 1) 95 21 ected for laborator isotopic 75 99 94 75 75 68 29 87 oned; S - Pb* 703 602 270 165 159 114 115 188 123 307 227 224 196 179 132 139 111 109 ppm oic r- tz diorite (T O6 O3 O1 oz U-Th-Pb SR SC EZ EZ EZ EZ EZ EZ EZ EZ EZ EZ EZ EZ EZ EZ EZ EZ EZ EZ EZ EZ EZ EZ ZC con Zi type^ . uhedrally z oter e1 centage of common 206Pb ain bl haean er ijala quar ijala paragneiss (T laeopr 4.1 1.1 7.1 6.1 2.1 3.1 8.1 9.1 5.1 1.1 3.2 7.1 9.1 6.1 spot 13.2 11.1 13.1 10.1 24.2 12.1 24.1 11.1 12.1 14.1 22.1 25.1 21.1 28.2 Gr Ta To To Arc Pa * Radiogenic Pb: corr # P ^ EZ - E Pre-1.91 Ga deformation and metamorphism in the Palaeoproterozoic Vammala Migmatite Belt… 113 8 9 5 6 8 5 9 9 8 4 4 5 7 7 7 7 7 9 7 ± 10 10 14 15 13 16 13 28 23 24 49 10 10 14 13 16 10 19 15 1942 1976 1994 1985 1983 1993 1993 1970 1977 2082 1975 1974 1932 1969 1989 2040 1926 2010 1973 1926 1975 1988 1978 2019 1970 1919 1977 1909 1966 1913 1924 1994 1984 1898 1908 1968 1931 1918 07/206 a) ±2 22 17 29 22 21 32 23 27 28 46 83 15 11 13 18 15 12 14 21 14 15 18 20 24 31 15 13 11 11 14 15 20 14 11 16 18 20 32 ent Ages (M ent Ages (M 1914 1983 2022 1957 1958 2067 1978 2012 2014 2151 1993 1825 1888 1995 1879 1932 1927 1908 1873 1944 1921 1984 2022 1927 1907 1972 1869 1922 1760 1882 1971 1913 1866 1947 1991 1933 1927 2018 06/238 par ppar Ap ±2 36 33 56 42 33 52 36 43 48 73 30 51 47 35 18 33 49 27 41 40 29 36 61 28 97 64 63 79 67 32 76 67 91 213 129 105 446 130 1913 1933 2105 1982 1966 2161 2056 2035 2023 2181 2002 1901 1850 2088 1922 1922 1947 1918 1882 2009 1848 1992 2042 1880 1304 1900 1788 1809 2054 1936 1016 1963 1843 2102 1602 2005 2219 2052 208/232 0.5 0.5 0.8 0.8 0.7 0.9 0.7 1.6 1.3 1.4 0.4 0.5 0.3 0.4 0.4 0.3 0.6 0.5 0.5 0.4 0.6 0.8 0.7 0.9 0.2 0.2 0.2 0.4 0.4 0.4 0.4 0.4 0.6 0.5 1.1 0.4 0.9 2.7 ±% Pb Pb*/ 206 207 0.11902 0.12132 0.12256 0.12195 0.12179 0.12250 0.12252 0.12093 0.12138 0.12883 0.12120 0.11840 0.12083 0.12220 0.12583 0.11799 0.12369 0.12110 0.11797 0.12130 0.12216 0.12149 0.12432 0.12096 0.11753 0.12137 0.11689 0.12068 0.11716 0.11785 0.12258 0.12191 0.11618 0.11679 0.12081 0.11834 0.11746 0.12125 % 1.4 1.0 1.7 1.3 1.3 1.8 1.4 1.6 1.6 2.5 0.9 0.7 0.8 1.1 0.9 0.7 0.9 1.2 0.9 0.9 1.1 1.2 1.4 1.9 0.9 0.8 0.7 0.7 0.9 0.9 1.2 0.8 0.7 1.0 1.0 1.2 2.0 4.8 U± Pb*/ 238 206 0.3457 0.3603 0.3685 0.3547 0.3548 0.3781 0.3592 0.3662 0.3668 0.3961 0.3623 0.3271 0.3402 0.3627 0.3385 0.3494 0.3483 0.3444 0.3372 0.3519 0.3471 0.3604 0.3684 0.3485 0.3443 0.3579 0.3364 0.3474 0.3140 0.3390 0.3577 0.3455 0.3357 0.3527 0.3619 0.3497 0.3483 0.3676 Finland 1.9 1.8 2.8 2.2 1.8 2.6 1.9 2.2 2.5 3.5 1.5 2.8 2.6 1.9 1.0 1.8 2.7 1.5 2.2 2.3 1.5 1.9 3.4 1.5 5.6 3.7 5.3 3.4 7.9 3.5 1.8 3.8 4.4 4.7 6.6 ±% 10.7 10.2 21.0 southern Th Pb*/ 232 208 0.0993 0.1003 0.1098 0.1031 0.1022 0.1128 0.1071 0.1059 0.1053 0.1139 0.1041 0.0986 0.0959 0.1088 0.0998 0.0997 0.1011 0.0995 0.0976 0.1045 0.0957 0.1036 0.1063 0.0975 0.0667 0.0986 0.0925 0.0936 0.1070 0.1005 0.0516 0.1020 0.0955 0.1096 0.0825 0.1043 0.1161 0.1068 belt, 1.2 1.4 2.0 1.6 1.0 1.5 0.9 1.3 1.7 1.9 0.3 2.5 2.5 1.6 0.6 1.5 2.1 1.1 1.9 1.8 0.8 0.9 2.5 1.2 5.7 3.4 5.3 3.2 7.8 3.3 1.5 3.6 4.2 4.5 3.7 ±% 10.6 10.1 20.7 migmatite Pb Pb*/ 206 208 0.17167 0.16408 0.16328 0.14521 0.20495 0.12583 0.26070 0.32409 0.19208 0.18045 0.26111 0.00561 0.02164 0.00312 0.04097 0.09982 0.03908 0.04796 0.10920 0.04908 0.03862 0.18402 0.26752 0.07386 0.00189 0.01445 0.00194 0.00790 0.00740 0.01310 0.00398 0.01086 0.02144 0.00873 0.02007 0.00805 0.00493 0.22332 mmala Va Pb the 206 0.073 0.016 0.114 0.056 0.054 0.132 0.165 0.395 0.261 0.451 0.038 0.003 0.003 0.041 0.169 0.033 0.005 0.009 0.065 0.049 0.016 0.017 0.042 0.130 0.050 0.008 0.010 0.015 0.068 0.016 0.033 0.002 0.023 0.017 0.173 0.043 0.086 0.558 % comm.# om fr Pb con Pb/ 204 206 zir 1.0E+05 1.4E+04 2.9E+04 2.9E+04 1.2E+04 9.7E+03 4.1E+03 6.1E+03 3.5E+03 4.2E+04 5.0E+05 5.0E+05 3.8E+04 9.4E+03 4.8E+04 3.3E+05 1.7E+05 2.4E+04 3.3E+04 1.0E+05 9.1E+04 3.8E+04 1.2E+04 3.2E+04 2.0E+05 1.7E+05 1.1E+05 2.3E+04 1.0E+05 4.8E+04 1.0E+06 7.1E+04 9.1E+04 9.3E+03 3.7E+04 1.9E+04 2.9E+03 2.2E+04 of U .01 .01 .01 Th/ 0.59 0.55 0.50 0.71 0.42 0.87 1.12 0.67 0.63 0.91 0.02 0.08 0.14 0.35 0.13 0.17 0.38 0.17 0.14 0.64 0.93 0.26 0.01 0.05 0.03 0.02 0.04 0.03 0.04 0.08 0.03 0.09 0.03 0.77 0.60 yses anal 70 90 30 96 79 61 58 43 13 53 99 73 58 47 38 35 18 71 33 20 30 19 23 45 14 43 13 28 Th 105 111 123 123 761 215 120 161 160 134 ppm pic isoto 73 72 34 U 87 69 36 179 175 158 156 145 141 110 838 719 694 690 710 615 543 347 319 282 252 906 685 672 636 600 502 487 464 225 224 ppm 1860 1350 1261 1110 52 61 81 61 27 32 42 41 25 89 20 14 00 02 71 71 61 63 59 60 50 36 31 76 16 86 Pb* 616 472 408 374 275 226 233 214 197 171 173 157 U-Th-Pb ppm r- S3 S2 S2 S2 O2 O2 O1 O1 O1 O1 O9 O1 O7 O4 (cont.) EZ EZ EZ EZ EZ EZ EZ EZ EZ EZ EZ LO LO LO LO LO LO LO LO LO LO LO LO LO con Zi type^ . e1 ain bl 2.2 3.1 2.1 8.1 4.1 5.1 spot 23.1 36.1 31.2 19.1 27.2 34.2 39.2 36.2 38.2 33.2 37.1 13.1 20.1 28.1 15.1 10.1 30.1 16.1 29.1 17.1 39.1 27.1 26.1 32.1 38.1 18.1 31.1 33.1 35.1 33.3 34.1 28.3 Gr Ta 114 Rutland,R.W.R.,Williams,I.S. and Korsman,K. ± 7 4 4 9 9 7 7 9 8 6 8 5 7 7 11 12 13 11 27 10 11 10 11 15 14 19 15 22 27 18 24 24 23 10 10 14 11 16 53 2704 2759 2906 2728 2923 2444 2698 2869 2668 2077 2757 1926 1953 2045 2010 2014 2023 1989 1989 1974 2006 2025 1965 2002 1982 1940 1987 2000 1935 1879 1890 1917 1977 1925 1908 1924 1982 1907 1872 07/206 a) ±2 34 44 31 34 27 17 30 14 29 21 33 15 15 20 18 20 19 27 26 24 35 37 51 33 38 13 13 29 42 11 12 11 21 15 22 19 26 22 40 ent Ages (M ent Ages (M 2520 2623 2826 2736 2913 2268 2507 2753 2554 2021 2690 1888 1883 1945 1967 1979 1973 1880 1978 1965 1964 2043 1921 1953 1920 1870 1949 1768 1865 1792 1763 1889 1940 1843 1879 1912 1947 1882 1889 06/238 par ppar Ap ±2 51 64 46 52 63 37 86 44 44 31 66 50 24 37 27 35 31 36 38 53 49 58 70 53 58 20 31 78 65 99 77 154 100 288 150 100 126 289 –– 550 846 2540 2622 2460 2820 2929 2257 2486 2638 2405 2036 2687 1962 1897 1908 1970 2017 1967 1845 1981 1971 1979 2069 1907 1963 1982 1877 1990 1924 2142 1902 1963 1302 2048 1819 1949 1779 208/232 ±% 0.4 0.7 0.7 0.8 0.7 0.3 1.6 0.2 0.6 0.5 0.7 0.5 0.4 0.4 0.5 0.6 0.6 0.8 0.8 1.1 0.8 1.2 1.5 1.0 1.3 0.4 0.4 1.3 1.2 0.4 0.3 0.4 0.4 0.6 0.6 0.8 0.6 0.9 2.9 Pb Pb*/ 206 207 0.18561 0.19196 0.21005 0.18835 0.21228 0.15891 0.18496 0.20533 0.18165 0.12843 0.19168 0.11802 0.11979 0.12614 0.12368 0.12399 0.12456 0.12225 0.12226 0.12120 0.12342 0.12472 0.12057 0.12313 0.12176 0.11890 0.12206 0.12300 0.11855 0.11493 0.11564 0.11742 0.12139 0.11792 0.11681 0.11789 0.12174 0.11676 0.11450 % 1.6 2.0 1.4 1.5 1.1 0.9 1.5 0.6 1.4 1.2 1.5 0.9 0.9 1.2 1.0 1.1 1.1 1.7 1.5 1.4 2.1 2.1 3.1 1.9 2.3 0.8 0.8 1.8 2.6 0.7 0.8 0.7 1.2 0.9 1.4 1.1 1.5 1.4 2.4 U± Pb*/ 238 206 0.4783 0.5021 0.5502 0.5288 0.5712 0.4217 0.4754 0.5328 0.4862 0.3682 0.5179 0.3404 0.3393 0.3523 0.3568 0.3593 0.3581 0.3385 0.3591 0.3563 0.3563 0.3730 0.3473 0.3539 0.3469 0.3366 0.3531 0.3156 0.3356 0.3204 0.3145 0.3404 0.3511 0.3309 0.3384 0.3452 0.3526 0.3390 0.3404 Finland ±% 2.2 2.6 2.0 2.0 2.3 1.7 3.7 1.8 2.0 1.6 2.6 2.7 1.3 2.0 1.4 1.8 1.7 2.1 2.0 2.8 2.6 3.0 3.8 2.8 3.1 1.1 1.6 4.2 7.5 5.5 5.2 4.5 6.7 12.0 11.9 15.3 11.9 16.8 southern Th –– Pb*/ 232 208 0.1339 0.1385 0.1294 0.1497 0.1559 0.1182 0.1309 0.1394 0.1264 0.1060 0.1422 0.1019 0.0984 0.0990 0.1023 0.1050 0.1022 0.0956 0.1030 0.1024 0.1029 0.1078 0.0989 0.1020 0.1030 0.0973 0.1035 0.0999 0.0276 0.0428 0.1118 0.0987 0.1020 0.0665 0.1067 0.0942 0.1012 0.0920 belt, ±% 1.0 1.1 1.2 0.9 1.1 1.4 3.3 0.7 1.1 0.8 1.9 2.4 0.7 1.3 0.7 1.3 1.0 0.7 1.1 2.3 0.9 1.7 1.3 1.6 1.5 0.7 1.4 3.6 7.3 5.3 4.9 4.2 6.5 11.5 11.7 15.3 11.8 16.7 migmatite Pb –– Pb*/ 206 208 0.19318 0.22704 0.26194 0.38332 0.28480 0.03047 0.06515 0.08577 0.10971 0.08378 0.15648 0.01963 0.12867 0.23532 0.19114 0.20123 0.19329 0.32314 0.21324 0.14867 0.28682 0.19698 0.35060 0.17683 0.16096 0.17940 0.06385 0.09668 0.02068 0.00256 0.00247 0.00377 0.00124 0.01052 0.01222 0.02309 0.01557 0.00665 mmala Va Pb the 206 0.120 0.068 0.168 0.116 0.228 0.026 0.018 0.057 0.105 0.065 0.112 0.033 0.101 0.076 0.018 0.032 0.106 0.080 0.093 0.115 0.042 0.149 0.286 0.137 0.024 0.025 0.048 0.354 0.291 0.079 0.027 0.014 0.010 0.394 0.036 0.097 0.189 0.170 0.708 % comm.# om fr Pb con Pb/ 204 206 zir 1.3E+04 2.3E+04 9.5E+03 1.4E+04 7.0E+03 6.3E+04 9.1E+04 2.8E+04 1.5E+04 2.5E+04 1.4E+04 4.8E+04 1.6E+04 2.1E+04 9.1E+04 5.0E+04 1.5E+04 2.0E+04 1.7E+04 1.4E+04 3.8E+04 1.1E+04 5.6E+03 1.2E+04 6.7E+04 6.3E+04 3.3E+04 4.5E+03 5.5E+03 2.0E+04 5.9E+04 1.1E+05 1.7E+05 4.1E+03 4.5E+04 1.6E+04 8.5E+03 9.4E+03 2.3E+03 of U .25 .01 .00 .02 .01 Th/ yses 0.69 0.82 1.11 1.35 1.04 0.11 0.24 0.33 0.42 0.29 0.57 0.07 0.44 0.84 0.67 0.69 0.68 1.14 0.74 0.52 0.99 0.68 1.23 0.61 0.54 0.62 0.22 0.31 0.02 0.01 0.05 0.04 0.08 0.05 anal 20 70 40 40 10 Th 65 76 36 72 71 63 38 48 25 11 29 13 29 12 222 135 160 138 104 133 171 120 186 223 372 222 179 174 267 151 124 127 405 210 ppm pic 72 isoto 34 56 U 62 70 44 321 164 144 102 961 560 523 284 639 548 503 445 332 260 257 234 203 140 125 104 103 103 652 963 882 837 549 346 346 221 158 ppm kaalinen 1) 1283 1033 01 21 88 14 86 09 54 45 47 73 46 98 84 55 54 44 45 41 27 76 52 14 Pb* 182 101 100 245 182 183 184 134 107 104 395 312 289 284 176 114 117 U-Th-Pb ppm oic r- kosic gneiss (I S2 O4 O2 O3 O1 O8 O3 O5 O2 oz (cont.) EZ EZ EZ EZ EZ EZ EZ EZ EZ EZ EZ EZ EZ EZ EZ EZ EZ EZ EZ EZ con LO LO LO LO LO LO LO LO LO LO Zi type^ . oter e1 ain bl haean laeopr aalinen ar 8.2 6.1 3.1 7.1 4.1 1.2 9.2 5.1 8.1 1.1 2.1 9.1 spot 23.1 24.2 17.1 11.3 18.1 22.1 28.1 20.1 16.1 14.2 26.1 29.2 13.1 12.1 15.1 21.1 25.1 19.1 27.1 30.1 11.2 11.1 24.1 10.1 29.1 31.1 14.1 Gr Ta Ik Arc Pa Pre-1.91 Ga deformation and metamorphism in the Palaeoproterozoic Vammala Migmatite Belt… 115

Fig. 10.Te ra-Wasserburg concordia diagram forzircon from the To ijala-1 paragneiss.

affected by Archaean metamorphism, were a zoned grain with 207Pb/206Pb ages of ~2.02 Ga present in the sediment’s source region. and the remainder at ~1.98 Ga. Grains with low The older Palaeoproterozoic group consists of (metamorphic) Th/U fall entirely within the most of the analyses of euhedrally zoned zir- younger sub-group (Figs. 10, 11). The ten anal- con, some of overgrowths and some of struc- yses of overgrowths in the younger sub-group, tureless zircon. The analyses aredispersed in including overgrowths with both low and mod- both 206Pb/238Uand radiogenic 207Pb/206Pb erate Th/U, give the same 207Pb/206Pb age with- (Fig. 10. Analyses of euhedrally zoned grains in analytical uncertainty, 1978±5 (tσ ) Ma. The arethe most dispersed, rangingin 207Pb/206Pb ages of the three structureless grains, two of apparent age from ~2.08 Ga to ~1.96 Ga,and which have low metamorphic Th/U, are indis- from concordant to ~15% discordant. Most tinguishable from this. (14 of 17), however, haveanapparent age of When all 27 analyses in the ~1.98 Ga group ~1.98 Ga.The radiogenic 207Pb/206Pb of all are combined, the radiogenic 207Pb/206Pb values these 14 is equal analytical uncertainty,giv- are found to scatter very slightly more than ex- ing aweighted mean age of 1981±6 (tσ )Ma. If pected from the analytical uncertainties. If the the four most discordant analyses areomitted four most discordant analyses are again omitted, the age calculated does not change significant- the scatter is eliminated, giving a weighted mean ly,1985±9 (tσ )Ma. age of 1979±4 (tσ ) Ma. This is the best estimate The analyses of overgrowths, one zoned grain of the age of the thermal event that produced and the structureless zircon are also dispersed in these various zircon types. both Pb/U and Pb isotopic composition. They Theyounger Palaeoproterozoicc group con- fall into two sub-groups, three overgrowths and sists of analyses of ten overgrowths, one struc- 116 Rutland,R.W.R.,Williams,I.S. and Korsman,K.

Fig. 11. Th/U vs age forthe To ijala-1 paragneiss and theIkaalinen arkosic gneiss as afunction of crystal zoning patterns. Note that the dashed reference ages do not all refertoinferredgeological events (see text).

tureless and three zoned grains. The structure- We havethereforeseparated the eight low less grain and most of the overgrowths havelow, Th/U overgrowths (Fig. 11) from the other metamorphic Th/U. Twoovergrowths havein- grain types. These overgrowths have equal 207Pb/ termediate Th/U ratios, and the four analyses of 206Pb within error, yielding a weighted mean age three zoned grains havehigher,igneous, values of 1916±5 (tσ ) Ma. This is considered the best (Fig. 11). Pb/U is dispersed, the analyses ranging estimate of the age of the youngest metamor- from concordant to 10% discordant, and there phic zircon growth in this rock. The remain- is aslightly larger range in radiogenic 207Pb/206Pb ing group of six analyses of disparate grain types than is expected from the analytical uncertain- gives a weighted mean 207Pb/206Pb age of 1927±8 ties. Some of the scatter is due to the radiogen- (tσ ) Ma (see discussion in section 7.2). ic 207Pb/206Pb of one high Th/U analysis being higher than the rest, but even the lowTh/U analyses alone aredispersed morethan expected. 5.3.Arkosic gneiss Ikaalinen-1 Thereisnoobvious outlier,however.The two The zircon from arkosic gneiss Ikaalinen-1 has overgrowths with higher Th/U also havehigh ra- broadly similar morphological and CL features diogenic 207Pb/206Pb,similar to that of the struc- to that from paragneiss Toijala-1, except that tureless grain and the three zoned grains. grains with relatively weak CL are more abun- Pre-1.91 Ga deformation and metamorphism in the Palaeoproterozoic Vammala Migmatite Belt… 117

Fig. 12.Te ra-Wasserburg concordia diagram forzircon from the Ikaalinen-1 arkosic gneiss. dant (Fig. 8). Analyses of 39 areas on 31 grains None of these overgrowths, however, has the are listed in Table 1 and plotted in Fig. 12. As very low Th/U commonly produced by meta- with the previous sample, structureless zircon morphism (Fig. 9), nor is their Th/U as high as and overgrowths, namely zircon morphologies that of the Archaean overgrowths from Toijala- that might record the late metamorphic history 1. The older analyses tend to become more dis- of the rock, were targeted preferentially. cordant with falling 207Pb/206Pb, raising the pos- The range of chemical compositions is very sibility that all or most are defining a discord- similar to that in the zircon from paragneiss Toi- ance line generated by the effects of an ear- jala-1; overgrowths and structureless grains have ly Pb-loss event on zircon of a single crystalli- on average higher U contents than the euhedral- sation age (Fig. 12). A line fitted to all analy- ly zoned grains. The analyses yield a wide range ses has an MSWD of 4. Without analysis 11.3 of apparent ages, ~2.92–1.87 Ga,with most the MSWD falls to 2, indicating near co-line- clustering in the range ~2.02–1.90 Ga.Most arity within analytical uncertainty. The concor- of the analyses are slightly discordant, particu- dia intersections, determined by a Monte Carlo larly those of the older grains, which introduc- line fit (Ludwig, 2001), are 2.98±0.04 Ga (the es added uncertainty in assessing the ages of zir- original zircon age) and 2.00±0.08 Ga (the pos- con growth. sible age of the Pb-loss event). An alternative in- Analyses of the older grains yield 207Pb/206Pb terpretation is to consider each small cluster of apparent ages in the range ~2.92–2.44 Ga. 207Pb/206Pb ages as representing a separate zircon Some come from zoned zircon, but several, par- age, in which case the older zircons would have ticularly those with lower 207Pb/206Pb, are anal- crystallisation ages of ~2.9, ~2.75 and ~2.69 Ga yses of overgrowths (grains 18 and 6 in Fig. 8). respectively. 118 Rutland,R.W.R.,Williams,I.S. and Korsman,K.

Fig. 13.Backscatteredelectron images showing different textures and ages of monazite from the To ijala-1 paragneiss and the Ikaalinen-1 arkosic gneiss (see text).

Fig. 14.Tera-Wasserburg concordia diagram formonazite from the To ijala- 1paragneiss and the Ikaalinen-1 arkosic gneiss.

As in Toijala-1, the Palaeoproterozoic zircon Omitting those, there is still excess scatter. Con- falls into two main age groups, one at ~1.98 Ga sidering only the zoned grains, the data are still and the other at ~1.91 Ga.Analyses of euhe- scattered, due to the analysis of one area with drally-zoned zircon have the largest range in ap- one of the highest Ucontents of that zircon parent age (~2.08–1.93 Ga) and lie mostly in type. Omitting this analysis also, the scatter in the older group (Fig. 12). The younger group the remaining eleven analyses is no longer sig- is dominated by analyses of low-Th/U over- nificant, yielding aweighted mean 207 Pb/206 Pb growths (Fig. 11). Analyses of structureless age of 2005±9 (tσ ) Ma. The other four analy- grains and higher-Th/U overgrowths aredis- ses in the group (all of overgrowths) give a mean tributed between both. Because most analyses 207Pb/206Pb age of 1983±14 (tσ ) Ma. are discordant, radiogenic 207Pb/206Pb provides The younger Palaeoproterozoic group consists the best means of determining the ages. of 12 analyses, most of which arefromover- The older Palaeoproterozoic group consists growths. The data areveryscattered, but di- of 18 analyses, two of which have significant- vide naturally into two age groups, the younger ly higher radiogenic 207 Pb/206 Pb than the rest. consisting of three analyses of low-Th/U over- Pre-1.91 Ga deformation and metamorphism in the Palaeoproterozoic Vammala Migmatite Belt… 119 growths (Fig. 8). The 9analyses in the older ameter), but also approximately equant, anhe- sub-group, including a structureless grain, two dral and pale amber coloured. They differ from euhedral grains, and overgrowths with both the Toijala-1 monazite grains in several ways, high and low Th/U, scatter more than expected however. Firstly, the grain surfaces are more ir- from the analytical uncertainties. Dividing the regular, with sharp cuspate projections and mul- sub-group in the same way as that from Toija- tiple scalloped embayments. Secondly, although la-1, the 5 oldest low-Th/U overgrowths yield a the grains show little fracturing in transmitted mean 207Pb/206Pb age of 1917±12 (tσ ) Ma and light,BSE imaging reveals some tobeclosely the three youngest an age of 1887±18 (tσ ) Ma, microfractured. Thirdly,BSE imaging shows the large uncertainty reflectingthe small number very little patchy zoning. Instead most grains of analyses pooled. The remaining zoned grains, are either structureless or broadly growth zoned structureless grain and high-Th/U overgrowth (Fig. 13), the truncation of the zones indicating probably have different origins (see section7.2) that they arethe remaining portions ofmuch but, if they are grouped, they have a mean 207Pb/ larger crystals. These features are consistent with 206Pb age of 1942±14 (tσ ) Ma. the Ikaalinen-1 monazite grains being the remnants of partly dissolved coarser monazite that 6. Monazite isotopic compositions has been deformed but not recrystallised. TenU-Th-Pb analyses of monazite from Small amounts of monazite were recovered from each sample are listed in Table 2. The monazite the two metasedimentary gneisses Toijala-1 and from Toijala-1 has a very wide range of U con- Ikaalinen-1. Ten grains were selected for analysis tent (190–4030 ppm) and Th/U (6.0–52.3). In from the former and 7 from the latter, all high contrast, the monazite from Ikaalinen-1 has a clarity and with minimal inclusions and micro- relatively uniform composition, U 4500–7300 fractures. ppm, Th/U 4.3–8.8, consistent with the relative The chosengrains from paragneiss Toijala- uniformity of its BSE response. The isotopic 1werefine (<80 µm diameter) approximate- analyses of both samples form tight clusters on ly equant, somewhat rounded, anhedral and a concordia diagram (Fig. 14), in both cases on pale amber coloured. Their surfaces arerel- average about 1.5% reverse discordant. The ra- atively smooth, although with some embay- diogenic 207Pb/206Pb in each sample is very con- ments consistent with confined growth. Inclu- stant, however, so the reverse discordance does sions and microfractures are rare. No zoning is not detract from the measurement of the ages. visible in transmitted light, but backscattered It could reflect late U loss from the monazite or electron (BSE) images show each grain to con- be due to the presence of initial 230Th, but as the sist of a patchwork of areas of different compo- 208 Pb/232Th apparent ages arealso higher than sition (Fig. 13). Tw o features suggest that mon- the 207 Pb/206Pb ages, it is more likely a Pb/U cal- azite with relatively strong BSE has been re- ibration error. placed, in some cases almost totally,bymon- The 207Pb/206Pb measurements from the To- azite with weak BSE. Firstly,monazite with ijala-1 monazite arethe moredispersed be- strong BSE commonly occurs towards the cen- cause of the lower Pb contents and hence larg- tres of the grains and in rare cases preserves weak er analytical uncertainties for that sample. All growth zoning. Secondly, the boundary between are equal within analytical uncertainty, howev- the monazite with strong BSE, and structureless er, yielding a weighted mean age of 1880.0±6.0 monazite with weak BSE is commonly convex (tσ ) Ma. The monazite from Ikaalinen-1 is sig- towards the former. nificantly older,with aweighted mean 207 Pb/ The monazite grains from arkosic gneiss 206Pb age of 1926.1±2.9 (tσ ) Ma. Ikaalinen-1 were slightly coarser (<100 µm di- 120 Rutland,R.W.R.,Williams,I.S. and Korsman,K.

Table2. U-Th-Pb isotopic analyses of monazite from the Vammala migmatite belt, southern Finland Grain Pb* U Th Th/ 206Pb/ % 206Pb 208Pb*/ 208Pb*/ 206Pb*/ spot ppm ppm ppm U 204Pb comm.# 206Pb ±% 232Th ±% 238U±% Toijala paragneiss (Toijala 1) 3.1 4709 4033 38183 9.5 1.4E+05 0.011 2.794 0.2 0.1019 3.0 0.3452 1.1 7.1 3116 3445 22798 6.6 1.4E+05 0.012 1.915 0.5 0.0999 3.5 0.3453 1.2 4.1 2598 1091 25442 23.3 3.7E+04 0.044 6.775 0.5 0.1012 3.6 0.3485 2.0 1.1 731 831 4948 6.0 9.1E+04 0.017 1.729 0.4 0.1039 3.9 0.3581 2.0 6.1 1332 802 11901 14.8 1.0E+06 0.001 4.289 1.6 0.1026 4.6 0.3550 3.7 5.1 2496 526 27339 51.9 1.0E+06 0.001 14.988 0.3 0.0980 3.5 0.3397 2.2 10.1 2397 490 25665 52.3 1.4E+05 0.012 15.286 0.6 0.1004 4.3 0.3437 2.8 9.1 1736 452 18072 40.0 7.7E+04 0.021 11.640 0.4 0.1012 3.4 0.3473 2.1 8.1 1148 334 12050 36.1 1.0E+06 0.002 10.321 0.9 0.0992 3.9 0.3466 2.8 2.1 470 187 4814 25.8 1.4E+05 0.011 7.442 0.6 0.0981 2.8 0.3395 1.9

Ikaalinen arkosic gneiss (Ikaalinen 1) 1.1 5354 7308 31416 4.3 7.7E+04 0.022 1.224 0.4 0.1031 2.8 0.3621 1.0 5.2 5787 6595 40576 6.2 1.0E+05 0.017 1.773 0.2 0.1012 3.1 0.3511 1.6 6.1 5553 6012 40065 6.7 4.8E+04 0.034 1.935 0.2 0.1016 2.4 0.3500 1.0 7.1 6020 5987 45976 7.7 2.8E+04 0.057 2.192 0.2 0.1003 2.8 0.3512 1.3 2.1 5661 5907 40709 6.9 4.8E+04 0.033 2.002 0.2 0.1032 3.0 0.3552 1.5 2.2 5641 5751 40045 7.0 2.0E+05 0.008 2.011 0.2 0.1047 3.0 0.3625 1.4 5.1 4817 5240 32980 6.3 2.3E+04 0.069 1.831 0.4 0.1049 3.0 0.3607 1.2 1.2 5445 4990 40953 8.2 3.0E+04 0.053 2.384 0.3 0.1047 3.0 0.3602 1.5 4.1 5237 4878 40807 8.4 6.3E+04 0.026 2.388 0.2 0.1011 2.8 0.3540 1.6 3.1 4936 4539 39857 8.8 3.2E+04 0.050 2.434 1.3 0.0981 3.3 0.3540 1.3 # Percentage of common 206Pb. * Radiogenic Pb: corrected for laboratory-derived surface common Pb using 204Pb.

7. Interpretation of the isotopic data TSB metamorphic zircon, then to distinguish it in turn from detrital metamorphic zircon de- Being metasediments, both Toijala-1 and Ikaa- rived from the metasediment sources. linen-1 contain zircon of a wide range of types, Monazite, on the other hand, is commonly compositions and ages. Most of this zircon is de- chemically unstable under low to medium grade trital, derived from a range of igneous, sedimen- metamorphic conditions (e.g. Kingsbury et al., tary and metamorphic protoliths. However, giv- 1993; Williams, 2001), and detrital monazite, en that the peak metamorphic grade of the two therefore, rarely survives to high grade, partic- metasediments is upper amphibolite to granulite ularly in metapelite. Monazite present in such facies, it is also to be expected that the samples rocks at upper amphibolite facies and above is will contain a small fraction of metamorphic zir- usually entirely metamorphic, retaining no iso- con grown in situ after deposition. If the thesis topic record of its former detrital age. is correct that the Vammala migmatite belt was Irrespectiveoftheir relation to the strati- metamorphosed to high grade before the wide- graphic record, three metamorphic events are spread igneous activity and high-grade meta- indicated by the present data, viz. morphism that followed deposition of the overlying volcano-sedimentary TSB, then it is pos- •~1.98 Ga,fromlow-Th/U overgrowths in sible that some metamorphic zircon from that both Toijala-1 and Ikaalinen-1. first post-depositional metamorphism might be •~1.92 Ga,fromlow-Th/U overgrowths in preserved. Toijala-1 and Ikaalinen-1 were select- both Toijala-1 and Ikaalinen-1, and from ed to be as free as possible from the effects of the monazite in Ikaalinen-1. post-TSB high-grade event, but nevertheless the •~1.88 Ga,fromlow-Th/U overgrowths in challenge remains to distinguish pre- from post- Ikaalinen-1, and from monazite in Toijala-1. Pre-1.91 Ga deformation and metamorphism in the Palaeoproterozoic Vammala Migmatite Belt… 121

207Pb*/ Appparparent Ages (Ma) 206Pb ±% 208/232 ±206/238 ±207/206 ±

0.11513 0.2 1961 56 1911 18 1882 4 0.11475 0.3 1925 65 1912 21 1876 5 0.11513 0.7 1948 67 1927 33 1882 14 0.11522 0.5 1999 75 1973 35 1883 9 0.11433 0.5 1974 88 1958 63 1869 10 0.11582 0.6 1890 63 1885 37 1893 12 0.11586 0.8 1934 79 1905 47 1893 15 0.11568 0.8 1948 63 1922 35 1891 14 0.11550 1.1 1912 71 1918 47 1888 19 0.11543 1.7 1891 52 1884 30 1887 31

0.11845 0.2 1983 54 1992 17 1933 3 0.11765 0.2 1948 58 1940 27 1921 4 0.11792 0.2 1956 45 1934 18 1925 4 0.11843 0.2 1931 52 1940 22 1933 4 0.11787 0.2 1985 57 1960 25 1924 4 0.11804 0.3 2013 57 1994 24 1927 6 0.11772 0.2 2016 58 1985 20 1922 4 0.11793 0.2 2012 57 1983 25 1925 4 0.11818 0.2 1946 51 1954 27 1929 4 0.11755 0.2 1892 60 1954 22 1919 4

The ~1.88 Ga event reflectsthe major post-TSB growths are necessarily younger than their host thermal episode, which has been recognised in grains, and must therefore be very close in age to much of the Svecofennian province.The two the metamorphic overgrowths. The number of older metamorphic events have not been recog- measurements is small, but it is noteworthy that nised hitherto. none of the overgrowths, whether low or intermediate in Th/U, has an Archaean host. This may indicate that Archaean grains were not 7.1.The ~1.98 Ga Event present at the site of metamorphism, but more Zoned grains of igneous origin dominate the measurements would be required to test this. zircon populations in both metasedimenta- The overgrowths in Ilkaalinen-1 are similar ry samples, but distinct overgrowths arerela- to those in Toijala-1 and four analyses (two with tively common. The analyses of ten low-Th/U lowand two with intermediate Th/U ratios) overgrowths in Toijala-1 identify ametamor- giveamean 207 Pb/206 Pb age of 1983±14 (tσ ) phic event at 1978±5 (tσ )Ma. Tw oofthree Ma. This indicates that the source areas of both structureless grains also have low Th/ U ratios samples were affected by the same early meta- and can be attributed to the same metamor- morphic event. However, while the main popu- phic event. lation of zoned grains for Tojala-1 yields an age The metamorphic age for Toijala-1 is not sig- of 1985±9 (tσ ) Ma, close to the age of metamor- nificantly different from the age of the main phism at 1978±5 (tσ ) Ma, the main population population of zoned grains at 1985±9 (tσ ) Ma. of zoned grains in Ikaalinen-1 yields an age of It is notable that a group of overgrowths has Th/ 2005±9 (tσ ) Ma. This is suggestive of rather dif- U ratios intermediate between the zoned grains ferent source areas for the two samples, but per- and the low-Th/U overgrowths. These over- haps from the same complex. Those source are- 122 Rutland,R.W.R.,Williams,I.S. and Korsman,K.

as were metamorphosed soon after the magma- as recording the post-depositional VD1 meta- tism and prior to erosion and deposition of the morphism. present metasediment. We speculate, on the ba- The evidence from the monazite in Ikaa- sis of the simple, coherent populations of zir- linen-1 confirms this interpretation. All grains cons in the two samples, that the metamor- yield the same age, 1926.1±2.9 Ma.Noold- phosed source complex for the VMB sediments er monazite is present, as would be expected if was also relatively nearby the sites of deposition: some detrital monazite were preserved. These more varied populations might be expected in rocks are known to have reached high grade (sil- far-travelled sediments. The possibility also aris- limanite-potassium feldspar,with fibrolite ages that the ~1.98 Ga metamorphic event was ac- gregates elongated parallel to S 0 /S1 planes [Kil- companied by tectonic activity that was respon- peläinen, 1998, p. 61]) during VD1 . Monazite sible for uplift and erosion, which led to the at such grades is usually entirely metamorphic deposition of the sedimentary protoliths of the and consequently this age can confidently be samples soon after the metamorphic event (Ta- attributed to post-depositional metamorphism ble 3). during VD1 . The age obtained shows that monazite re- crystallisation could not take place at this lo- 7.2.The ~1.92 Ga Event cality after about 1.92 Ga (as it did elsewhere 7.2.1. Correlation with VD1 metamorphism in the VMB). This probably implies achange of conditions with uplift and erosion after

Asub-set of eight low-Th/U overgrowths in VD1 .Asdescribed in section 6, the monazite the younger Palaeoproterozoic group from To- appears to consist of remnants of partly dis- ijala-1 gives a metamorphic age of 1916±5 (tσ ) solved coarser monazite that has been micro- Ma.The analogous sub-set of fivelow-Th/U fractured in alater event, but not recrystallised. overgrowths in Ikaalinen-1 gives a mean age of We consider that this later event can probably 1917±12 (tσ )Ma. Thewell-dated host grains be correlated with the retrogressivemuscovite to the overgrowths aremuch older.Some are that occurs in this area and in other areas im- of similar age to the main population of zoned mediately adjacent to the TSB(Fig. 5). It can grains in the older Palaeoproterozoic groups and be attributed to retrograde metamorphism that others are of Archaean age (Fig. 8). Thus this may haveoccurred at any time after ~1.92 Ga, metamorphic event was superimposed on zir- but most probably during the subsequent met- con from diverse, and significantly older sourc- amorphism of the TSB. This interpretation is es. As discussed below, deposition of the meta- supported by previously unpublished evidence sediment protolith probably immediately post- from asmall gabbrobody intruded into silli- dated the ~1.98 Ga metamorphism. manite gneiss about 7km SSE of Ikaalinen-1 Both metasediment samples were selected and about 2km SW of Herttuala, in the SE from locations displaying only the VD1 schis- corner of the 1:100 000 Ikaalinen map area tosity, and the field evidence for Toijala-1 shows (Virransalo &Vaarma, 1993; GR, 6831.55N, that VS1 is older than the nearby granitoid, dat- 2453.10E). The gabbroclearly post-dates the ed at 1888±5 Ma.Both areas were metamor- sillimanite metamorphism and the oldest VD1 phosed to high grade during VD1 , which should schistosity.Itisconsidered to represent afeed- have resulted in further zircon growth. The low- erchannel of the Tamperevolcanic rocks (P. Th/U zircon overgrowth ages at ~1.92 Ga are Peltonen, pers. comm., 1992), and is itself de- the only possibility for this VD1 metamorphic formed and metamorphosed, presumably dur- event. We therefore interpret the zircon meta- ing the post-TSB TD1 event. Thus the preser- morphic ages of 1916±5 Ma and 1917±12 Ma vation of the VD1 monazite can be attributed Pre-1.91 Ga deformation and metamorphism in the Palaeoproterozoic Vammala Migmatite Belt… 123 to the locality b eing at arelative ly high struc- The structureless grain 13 in Toijala-1 (Fig. tural level during later deformation, and under 8) has alow Th/U ratio and gives an age of retrograde conditions such that monazite again 1932±9 Ma. As with the structureless grains as- began to dissolvebut was not recrystallised. sociated with the ~1.98 Ga metamorphism, this The monazite age is indistinguishable from grain can be attributed to growth during the the age of the low-Th/U zircon overgrowths same metamorphic event as produced the low- in the same sample. We consider that it is un- Th/U overgrowths.Similarly,the overgrowths doubtedly post-depositional and that it con- 28.1 and 10.1, with intermediate Th/U ratios firms that the metamorphic age recorded in and ages of 1926±5 Ma and 1926±9 Ma respec- both zircon samples was also post-depositional tively, can also be attributed to growth closely and correlates with the observed VD1 metamor- associated with the post-depositional metamorphism. It is of course possible that the metamorphism. Those zoned grains with similar appar- phic episode occupied a considerable time inter- ent ages (e.g. 12.1 and 9.1) are U-rich and prob- val and that the zircon and monazite are record- ably isotopically disturbed. Such disturbance is ing slightly different times within that interval. evident in the disparity between the ages meas- However, we consider that the metamorphism is ured on the core of grain 28. best characterised from the combined monazite The ages measured on the analogous sub- and zircon data with an age of ~1.92 Ga. group in Ikaalinen-1 can be similarly explained. One structureless grain (19, Fig. 8) gives a similar discordant age within error (1940±8 Ma) to 7.2.2. Age of deposition of the VMB sediments grain 13 in Toijala-1, but has a high Th/U ra- Lahtinen et al. (2002, p. 100) have noted that tio. The core of grain14.2 (Fig. 8) is very weakly evaluation of maximum deposition age should zoned, but has a very low Th/U ratio and yields bebased only on concordant analyses, prefer- 1926±9 Ma. Overgrowth analysis 11.2 with in- ably of several grains. This assumes that none termediate Th/U is strongly discordant and of those grains is produced or affected by post- gives an age of 1935±23 Ma. depositional metamorphism. Our analyses have We therefore interpret the analyses of over- been particularly directed at identifying post- growths and structureless grains in both samples depositional growth, so some types of zircon as representing post-depositional growth during havebeen analysed that would normally have the onset of the ~1.92 Ga metamorphism, anal- been avoided. The challenge is to distinguish ogous to the similar grains seen in association concordant pre-depositional zircon from post- with the ~1.98 Ga metamorphism. The few as- depositional growth and isotopic resetting (dis- sociated ages from zoned grains, mainly discord- cordance). ant and high U, we interpret as affected by iso- The younger Palaeoproterozoic group in both topic resetting. As they include no analyses of samples is dominated by the low-Th/U over- normal well-zoned concordant igneous grains, growths produced by post-depositional meta- we consider that they do not bear on the max- morphism. The groups also include afew atyp- imum age of deposition. The sediment must ical structureless or poorly zoned grains, and have been deposited after ~1.98 Ga, the age of overgrowths with higher Th/U, but no unequiv- the youngest group of igneous detrital cores. It ocally igneous grains. The disparate grains evi- must have been deposited before the ~1.92 Ga dently havearange of origins, and as noted in post-depositional metamorphism. section 5above,they areonaverage slightly old- We have noted abovethatthe more near- er (1927±8 Ma for Toijala-1, and 1942±14 Ma ly concordant zoned grains in the older Pro- for Ikaalinen-1) than the low-Th/U overgrowths terozoic populations form groups older than (1916±5 Ma and 1917±12 Ma respectively). the ~1.98 Ga metamorphism. The lowest age 124 Rutland,R.W.R.,Williams,I.S. and Korsman,K. in the total Toijala-1 population is 1964±6 imum deposition a ges of “about 1.92-1.91 from grain 11.1, but this analysis shows high Ga”(Lahtinen et al., 1992, p. 100), and to imply U, is strongly discordant and also has an un- that deposition “took place c. 1.9 Ga shortly be- usually lowTh/U ratio.Pbloss is areal possi- fore the felsic volcanism” (Huhma et al., 1991, bility.Thereare no concordant zoned igneous p. 176; cf.Nironen, 2004; Kähkönen et al., grains definitely younger than the ~1.98 Ga 2004). While the uncertainties in the available metamorphism. Thus the VMB protoliths do data would allow an older maximum deposition notappear to havesampled any zircon-bear- age, we consider that the published interpreta- ing source areas aged 1.97–1.93 Ga,and they tions are reasonable. They are supported by (1) were probably deposited soon after the ~1.98 the conformable and interfingering relationship Ga metamorphism. Significantly,deposition of between the greywackes and the dated volcanics the sediments after ~1.97 Ga would allowthe (2) the presence in these rocks of only the same associated mafic rocks in the south of the VMB post-TSB deformation as in the overlying vol- to represent an extensional episode of similar canic sequences and, conversely, (3) the absence age to the 1.95–1.96 Ga Outokumpu and Jor- of the complex folding and of the VD1 schis- mua ophiolites. tosity, which are present throughout the VMB. In summary therefore both samples indicate Most importantly, however, it can now be add- a maximum deposition age of ~1.97 Ga, a min- ed that, in the SE Ikaalinen area from which imum of ~1.93 Ga, and post-depositional met- sample Ikaalinen-1 was taken, the VMB actual- amorphism at ~1.92 Ga.Thus the protoliths ly underlies the TSB succession with discordant of the VMB aremarkedly different from the contacts (section 2.3 above). Thus, the date of older sediments of the Bothnian Basin, which ~1.92 Ga obtained for the VD1 metamorphism have been inferred to have a maximum depo- there itself provides a maximum deposition age sition age of ~1.91-1.92 Ga and which com- for the overlying TSB sediments. The previously monly contain significantdetritalpopulations inferred maximum deposition age for the Myl- (measured on near concordant igneous grains) lyniemi Formation is just what would be expect- younger than ~1.95 Ga (Claesson et al., 1993, ed if the greywackes of that formation were the Lahtinen et al., 2002, both including discussion result of uplift and erosion following the crustal of sample A1, the location of which is shown on shortening associated with the VD1 event. The Fig. 3, and also sample A57 in the TSB). upward-fining trend in these sediments accords well with this interpretation. The Myllyniemi formation has been corre- 7.2.3. Relation of the ~1.92 Ga metamorphism to the lated with the Osara formation, a smaller area TSB and Haveri sequences of metaturbidites and migmatitic metatur-

The VD1 metamorphic episode at ~1.92 Ga is bidites to the NW of the main synclinal depres- older than the whole of the TSBvolcanic se- sion (Fig. 3; Kähkönen, 1999, p. 17). The Osa- quence, which has yielded a maximum age of ra formation appears to be underlain conform- 1904±4 Ma near the base (Kähkönen et al., ably by the basalts of the Haveri Formation, re- 1989). This in itself invalidates the earlier con- garded as the oldest unit in the TSB (Kähkönen clusion that the “migmatite area was merely a & Nironen, 1994). The latter authors showed marine partofthe Tampereisland arcforma- that these basalts have marginal basin affinities tion” (Kilpeläinen et al., 1994, p. 27). and suggested (op. cit., pp. 156–157) that they The youngest detrital zircons (e.g. 1907±15 might represent an “ensialic, relatively small, ex- Ma, Huhma et al., 199l) from the greywackes of tensional basin, floored by attenuated (Palaeo- the Myllyniemi formation underlying the vol- proterozoic) continental crust.” They (op. cit., canics havebeen considered to indicate max- p. 158) did not favour a protracted evolution- Pre-1.91 Ga deformation and metamorphism in the Palaeoproterozoic Vammala Migmatite Belt… 125 ary history, and considered that “tholeiitic ma- resent an extensional episode in t he evolution fic volcanism in an extensional setting was fol- of the Svionian Basin and the correlation with lowed, possibly continuously, by arc-type mag- the above ophiolites, canvassed by Vaasjoki & matism”. Huhma (e.g. 1999, p. 151) becomes geological- The ages of the Osara and Haveri formations ly plausible. Further work would be required to are not well established, however. The contact test this possibility, but we tentatively interpret with the overlying Harhala Formation is not ex- the Haveri and Osara formations to be part of posed, but in accordance with their viewthat the pre-1.92 Ga VMB. If this is correct it rein- the Haveri and Osara formations are closely as- forces the conclusion that the true TSB succes- sociated in time with the overlying arc volcan- sion is younger than ~1.92 Ga. ics, Kähkönen & Nironen, (1994, p. 146) cor- It follows from this conclusion and the dat- related the earliest schistosity in the Haveri dis- ing of VD1 that the widely accepted strati- trict with TS1 in the central TSB. If this is cor- graphic correlation between the t urbidites of rect, the Osara and Haveri formations may well the TSB(the Myllyniemi formation) and the be younger than ~1.91 Ga,but the structural metasediments in the north of the VMB can- pattern is relatively complex. The three defor- not be sustained. This correlation was large- mation episodes identified (op. cit.) are compa- ly based on lithological and geochemical sim- rable in style with the three episodes recognised ilarities (Lahtinen, 1996), and i tcan be sug- in the VMB (Kilpeläinen, 1998). S 1 is layer par- gested instead that these similarities are a con- allel and apparently not related to upright fold- sequence of the derivation of the TSB turbidites ing. Migmatitic mica gneisses occur adjacent to from the erosion of the upper parts of the VMB the overlying Harhala formation of the TSB, complex. and the metamorphism has affinities with that of the SW Ikaalinen area of the VMB (Fig. 5). 7.3.The ~1.88 Ga Event. The main folding is F 2 , while open F 3 folds are sub-parallel to the contact with the Harhala for- The monazite from Toijala-1 records a thermal mation. We suggest therefore that F 3 correlates event at 1880.0±6.0 Ma.Asdescribed above, with TS1 in the TSB and that these rocks repre- this monazite is chemically variable and textur- sent a relatively low-grade part of the VMB. ally different from that in Ikaalinen-1, consist- Vaasjoki & Huhma (1999) observed that the ent with almost total replacement of an earli- least radiogenic leads at Haveri and in the Ou- er monazite. It possibly replaced an older meta- tokumpu ophiolite complex are similar “and the morphic monazite that was formed during the two occurrences can be coeval”. Their lead iso- earlier ~1.92 Ga VD 1 metamorphism recorded tope data yields “rather uncertain age estimates in the zircon data, and in the Ikaalinen-1 mon- in the 1900-2000 Ma range”,which arecon- azite data. The possibly original monazite (e.g. sistent with the possible correlation but which 5.1, Fig. 8), however, appears to preserve no iso- they observe cannot be treated as true ages. We topic record of that event. have suggested above (section 7.2) that our sam- The ~1.88 Ga event is not recorded in the pled VMB sediments may have been deposited zircon from Toijala-1. The sample was selected as early as ~1.97 Ga and that such a deposition- from a location where the early VD1 schistosi- al age would allowthe associated mafic rocks ty was dominant. Younger leucosome veins are in the VMB proper to be of similar age to the present as a conjugate network crosscutting the 1.96–1.95 Ga Outokumpu and Jormua ophio- schistosity, but were avoided in sampling. Per- lites. It is possible therefore that the Haveri ba- haps the ~1.88 Ga event did not produce new salts could also be of this age, consistent with zircon growth because fluids,associated with the the Pb isotope data. They would then also rep- veins, were unable to penetrate the rather fine- 126 Rutland,R.W.R.,Williams,I.S. and Korsman,K. grained schistosity, in the absence of coeval pen- 7.4. Relationship of ages to evolution of meta- etrative deformation. morphism and migmatisation A weak indication of the ~1.88 Ga event is In their study of migmatites in the VMB, Mou- present in Ikaalinen-1 zircons. The three young- ri et al. (1999) found agreement in age within est Th/U overgrowths form aseparate group sometimes large error margins between U-Pb with amean age of 1887±18 Ma,although dating of monazite, limited SIMS dating of un- two of the analyses are discordant and the other zoned overgrowths on zircon, and Sm-Nd gar- has a large analytical uncertainty. One of these net-whole rock analyses. They suggested “that overgrowths (11.1) is superimposed on an old- the migmatites were metamorphosed at ca. 1880 er overgrowth (11.2) of the ~1.92 Ga metamor- Ma during asingle metamorphic event. Since there phic generation (Fig. 8), and thereforeclearly is no evidence of an earlier event in the study area, it records ayounger event. If these overgrowths is assumed that 1880 Ma is the approximate age of do represent the ~1.88 Ga event, they can prob- the peak metamorphic event.” However, only the ably be attributed to fluid access to this coarse- monazite yielded aprecise age, and therenow is grained metasediment at this time, associated evidence of an earlier event: the various meta- with TD1 migmatisation around major intru- morphic assemblages in the region clearly span sions to the west. a40m.y.period between the metamorphism at As outlined in section 2.1, an age of ~1.88 Ga ~1.92 Ga in the VMB and the peak metamor- can be inferred for emplacement of plutons and phism and deformation in the TSBat~1.88 Ga. the TD1 folding and metamorphism in the TSB Afull re-evaluation of that metamorphic history (Nironen, 1989b). This age correlates closely is beyond the scope of this paper but some sug- with the thermal event age of 1880.0±6.0 Ma gestions can be made to reconcile the newevi- obtained from the Toijala-1 monazite in the dence with that of Mouri et al. (1999). area of the VMB characterised by the garnet We havechosen VMB samples in which only sillimanite association (Fig. 5). Monazites dat- the earliest schistosity (VD 1 ), parallel to litholog- ed by Mouri et al. (1999) from mesosomes and ical layering, is present, and wherethe influence leucosomes in VMB migmatites at Luopioin- of younger events is apparently m inimal. This en, about 20 km ESE of Tampere, gave a similar has allowedthe identification of apost-deposi- mean age of 1878.5±1.5 Ma. These migmatites tional metamorphic event at ~1.92 Ga and its are characterised by the sillimanite + cordierite correlation with the metamorphism accompa-

± garnet association, which is largely restricted nying VD1 .Sillimanite and garnet both formed to the Tampere map sheet (Fig. 5), and probably early (Kilpeläinen, 1998, pp.96–97), and are represents a higher structural level. considered to be coeval with VD1 ,asare the thin We infer that the close correspondence in bands of concordant leucosome. The local field monazite ages between the TSBand VMB evidence shows that the VD1 metamorphism oc- events at different structural levels implies that curred beforethe emplacement of the Toijala- uplift and cooling took place rather rapidly after 2granitoid at 1888±5 Ma;and both the geo- the deformation of the TSB. As noted in section chronology and the broader structural evidence 2.1, the south-side-up shear zone, which sepa- discussed above suggest that the earliest schis- rates the TSB from the VMB (Fig. 3), is essen- tosity was formed beforethe TSBsuccession was tially a TD1 structure (Nironen, 1989a, p. 31) deposited. The existence of an early (Svionian) and could account for rapid decompression and metamorphism is thereforewell established. cooling as uplift and erosion of the VMB oc- In contrast, Mouri et al. (1999) selected sam- curred. We therefore conclude that the monazite ples to study the migmatisation that is superim- ages of ~1.88 Ga in the VMB correspond closely posed on the earliest schistosity and their data to the age of peak metamorphism in the TSB. are dominated by the younger events. It is likely Pre-1.91 Ga deformation and metamorphism in the Palaeoproterozoic Vammala Migmatite Belt… 127

Table3. Summaryofisotopic results and interpretations Zircon Tooijala-1ijala-1 Ikaalinen-1 Pre-depositional Archaean Fewer at ~2.72 Ga More at ~2.92–2.44 Ga. Abbsencesence of 2.4–2.1 Ga detrital grains Pre-depositional older Palaeoproter- Igneous complex Igneous complex ozoic (Within or on margin of older-stage (Within or on margin of older-stage Svionian basin?) at ~1.99 Ga Svionian basin?) at ~2.01 Ga Meetamorphismtamorphism at ~1.98 Ga Meetamorphismtamorphism at ~1.98 Ga Absence of post-1.97 Ga detrital igneous grains Erosion and deposition of younger-stage Svionian Basin soon after ~1.98 Ga? Post-depositional Palaeoproterozoic Early metamorphic? at ~1.93 Ga Early metamorphic? at ~1.94 Ga

Meetamorphismtamorphism and VD1 at ~1.92 Ga Meetamorphismtamorphism and VD1 at ~1.92 Ga Monazite Meetamorphismtamorphism and VD1 at ~1.93 Ga Unconformity above VMB

Deposition of TSB succession between VD1 and TD1 .

Intrusion of post-VD1 Tooijala-2ijala-2 granitoid at ~1.89 Ga Monazite Metamorphism and TD1 at ~1.88 Ga in the regional context that the earliest schistosi- ic events. Some weakly oscillatory-zoned rims ty observed was initiated as VD1 , but this is not were considered to be of magmatic origin and certain, as the schistosity was not demonstrated two of these in the L1 leucosome gaveages of to be parallel to lithological layering. Similarly, 1904±17 Ma and 1915±9 Ma.Afurther meta- it is not certain that the earliest leucosome vein morphic overgrowth in the mesosome gavean

L1 studied by Mouri et al. was coeval with VD1 . age of 1907±8 Ma. These were discounted as In any case, the mesosome assemblage was clear- possibly mixed ages by Mouri et al., but it now ly modified by younger events (op. cit., p. 34), seems possible that they representthe early (Svi- most notably the development of cordierite co- onian) metamorphism and melt formation. It is ronas around garnet. Cordierite does not occur also notable that IDTIMS zircon ages indicat- in the earliest concordant leucosome, L1, nor in ed for the L1 and L2 leucosomes were strikingly the discordant L2 leucosome, but it does occur different at 1950 Ma and 1892 Ma,and, if the in rare small L3 leucosome patches overprinting ages reflect mixed age populations (cf.op. cit., the main foliation. p. 47), these were evidently very different. Since Mouri et al. (op. cit., p. 38) consider that ‘the the peak metamorphic event affected both the textures observed in the mesosome and the for- leucosomes (op.cit., pp.47and 51), it apparent- mation of the different types of leucosomes can ly correlates with the L3 event or later.Wethere- be explained by progressive partial melting re- foreconsider that the data presented by Mouri actions during increasing temperature and de- et al. (1999)allows an extended metamorphic compression.’ They evidently assume that par- historybeforetheir peak metamorphic event. tial melting and the formation of the different At deeper levels in the migmatites, recrystal- leucosome generations occurred very shortly be- lisation of monazite occurred during the ~1.88 fore the peak metamorphic event but the availa- Ga metamorphic episode to produce the mon- ble evidence does not require this. azite of Toijala-1 and of the samples studied by Mouri et al. (op. cit., pp. 39–40) consider Mouri et al. (op. cit.) at ~1.88 Ga. We have in- that both the L1 and the L2 leucosome melts ferred above that this age corresponds closely to were produced alongwith garnet beforethe ap- theage of peak metamorphism (Bothnian) in pearance of cordierite in the assemblages. How- the TSB, which accompanied the post-deposi- ever no zircons, among the limited number of tional folding. The age of 1888±5 Ma obtained SIMS measurements, were specifically iden- from the Toijala-2 quartz diorite is very similar tified as being associated with these magmat- to the monazite ages, and also to the age of vol- 128 Rutland,R.W.R.,Williams,I.S. and Korsman,K. canics in the upper part of the TSB succession atite evolution, when the temperatureprobably (1889±5 Ma, Kähkönen et al., 1989), and to the remained above 700°C (op.cit., p. 41). age of granitoid clasts in that succession (Niro- Some granitoid emplacement accompa- nen, 1989b). The geological evidence there- nied theupright TD 1 folding and peak meta- fore shows that granitoid emplacement, uplift, morphisminthe TSB(Nironen, 1989a). The erosion and conglomerate deposition,all took youngest L-granite of Mouri et al. (op. cit., pp. place during the later part of volcanic activity 36 and 46) was probably intruded at this stage. in the TSB at ~1.89 Ga. The granitoids in the The granite is sillimanite-bearing and it pre- VMB are closely associated with migmatisation serves the same monazite age as the geological- in their contact zones and it therefore appears ly older samples. This confirms the correlation that the younger (Bothnian) metamorphic epi- of the monazite ages with the post-TSB meta- sode in the VMB began at least as early as ~1.89 morphism. Ga. It may have begun earlier at ~1.90 Ga, as The principal conclusions from Section 7 indicated by the early felsic volcanism dated at above are summarised in Table 3, which also in- 1904±4 Ma in the overlying TSB. We therefore corporates some conclusions from the following suggest that in the deeper levels of the VMB, re- discussion. newed metamorphism and migmatisation probably occurred during an extended period be- 8. Discussion: implications for Sve- tween ~1.90 and 1.88 Ma. cofennian tectonics In this geological context, the discordant L2 leucosomes were probably generated in as- The most important conclusions from our data sociation with the major granitoid intrusions and interpretations are: at ~1.89 Ga. Formation of the concordant L1 1) The VMB has a pre-1.92 Ga history of sed- leucosome may have occurred during the early imentarybasin development probably ex- stages of the renewed metamorphism, or it may tending back to ~1.97 Ga,which has not have occurred during Svionian metamorphism previously been recognised. The detrital zir- as indicated by the limited SIMS data noted con population provides evidence of apre- above. The latter would require an hiatus before depositional metamorphic episode at ~1.98 the discordant L2 leucosomes were generated. Ga and of ~2.1–1.98 Ga source complexes. We suggest that the later,post- L2, stages of 2) A major episode of deformation and meta- evolution involving cordierite growth after gar- morphism occurred at ~1.92 Ga before the net, and the production of cordierite plus melt deposition of the TSB succession. (Mouri et al., 1999, p. 39), were probably con- 3) The TSBsuccession was deposited on the tingent on the decompression of the subjacent VMB basement in the 40 m.y. interval be- VMB implied by the uplift and erosion at ~1.89 tween the tectonothermal episodes at ~1.92 Ga recorded in the TSB. This implies that the Ga and ~1.88 Ga. formation of the L1 and L2 leucosomes took Each of these conclusions (see also Table 3) bears place beforethis uplift. On the other hand, the on the interpretation of the tectonic relationship monazite ages of ~1.88 Ga in the VMB corre- of the VMB to adjacent tectonic domains, and spond closely to the age of peak metamorphism also has wider implications for the evolution of in the TSB, which followedthe deposition of the Svecofennian Province as a whole. the whole TSBsuccession. We thereforesuggest that, in order to givethe ages obtained, equilibra- tion of the Sm-Nd garnet-whole rock and U-Pb 8.1. Source of detrital zircons and two-stage ba- monazite systems studied by Mouri et al. must sin evolution haveoccurred during these later stages of migm- Previous studies of detrital zircon in sediments Pre-1.91 Ga deformation and metamorphism in the Palaeoproterozoic Vammala Migmatite Belt… 129 of the Bothnian B asin (Huhma et al., 1991; of Rutland et al. (2001b,Fig. 3an dp.232) Claesson et al., 1993; Welin et al., 1993) re- may reflect the ~1.98 Ga event, and the young- vealed that 2.1–1.9 Ga source areas were dom- er stage of Svionian Basin development may be inant, but the location of these was problemat- absent in that area. An older stage basin also ap- ic (e.g. Kähkönen, 1999, p. 27). Possible source pears to be present in the basement of the East areas of this age are not known in the adjacent European Platform, ~500 km to the SSE of the Karelian Province in Finland. The present study Tampere district (Claesson et al., 2001), where is the first test of the hypothesis (Rutland et rocks 2.10–1.95 Ga old are known from drill al., 2001b, p. 256) that the postulated pre-1.9 hole samples. The concealed NE-trending Ar- Ga metamorphic complexes were the principal chaean-Palaeoproterozoic boundary there (Fig. source, and does show that the VMB rocks con- 1) appears to havequite different characteris- tain appropriate zircon populations. tics (e.g. Bogdanova et al., 1996; Claesson et al., As suggested in section 7.1 the source com- 2001) from the NW trending boundary in Fin- plexes of the VMB sediments were probably land. It is marked by a wide ~2.0–2.1 Ga volca- not far distant from the sites of deposition. It no-plutonic belt (the Osnitsk-Mikashevichi Ig- is therefore possible that these source complex- neous Belt, OMB) and by ~2.1 Ga granulite es were themselves located within the Svioni- facies metamorphism in metasediments overly- an marginal basin. Possibly they were located ing the Archaean craton of Sarmatia to the SW. in island arcs separating the basin into a com- The trends of the Palaeoproterozoic belts to plex of back-arc and intra arc basins. The recog- the NW of Sarmatia are roughly parallel to the nition that the source complexes for the VMB cratonic margin in the south, but bend north- sediments were metamorphosed at ~1.98 Ga westwards or even westwards approaching the also suggests that theremay have been simul- Gulf of Finland (Fig. 1). This Palaeoproterozo- taneous tectonic activity which might separate ic province is apparently dominated by metased- basin evolution into older and younger stages, iments that contain detritus significantly older and which might correspond to accretion of an thanthe orogenic deformation and metamor- older (~2.1–1.98 Ga) stage. As suggested above phism.The range of initial ε Nd values is sim- (section 7.2), the associated mafic rocks in the ilar to that found in Svecofennian metasedi- younger stage basin (~1.97–1.92 Ga), notably ments (Claesson et al., 2001, p. 14), and these in the south of the VMB, and possibly at Ha- sediments can reasonably be interpreted as mar- veri, could represent a renewed phase of exten- ginal basin fill. In the context of a marginal ba- sion that can be tentatively correlated with the sin interpretation, the OMB can then be com- Outokumpu and Jormua ophiolites. Thus the pared with the Cenozoic Sikhote–Alin Belt on proposed two stages of evolution of the Svionian the margin of the Chinese craton, inboard of the Basin appear to correspond broadly in age (but Japan Sea marginal basin (e.g. Okamura et al., not in detrital zircon content) to the Lower and 1998), rather than with an Andean belt. Upper Kalevian tectofacies on the Karelian cra- An older Palaeoproterozoic domain (the ton margin (Kohonen, 1995, pp. 62–68), which Vitebsk Granulite Domain) adjacent to the cra- are also separated in age by the ophiolites. ton displays was metamorphosed to granulite The presence of granitoids as old as 1.96 Ga, facies at ~1.96 Ga, but is considered to have a intruding older, previously deformed, sequenc- crust formation age of ~2.0 Ga (Claesson et al., es in some parts of the Svecofennian in Sweden 2001, p.15). It could therefore represent an ear- (Wasström, 1993; 1996; Claesson & Lundqvist, ly phase of marginal basin development that was 1995) gives some direct support to the existence accreted by 1.96 Ga, and correspond to our pro- of the proposed older stage basin. Indeed, the posed older stage Svionian Basin. It is presuma- early deformation of the Storuman sub-zone bly rich in detritus from the ~2.1–2.0 Ga OMB 130 Rutland,R.W.R.,Williams,I.S. and Korsman,K. volcano-plutonic belt. This basin may have been er than ~1.90 Ga. We shall refer to this wide- bounded to the west by the ~1.98 Ga Okolovo spread deformation as Early Svecofennian (DE ), volcanic belt (Fig. 1), which has yielded U-Pb acknowledging that it may be somewhat dia- ages of 1982±26 Ma and 1998±10 Ma from chronous. Daly et al. (2001) show that the prin- juvenile metadacites (similar to the ages of the cipal deformation and metamorphism associ- main detrital populations in our Toijala-1 and ated with accretion inthe Lapland-Kola Oro- Ikaalinen-1 samples from the VMB). A meta- gen also took place before ~1.91 Ga. The age of morphic age of ~1.92 Ga (op. cit., p. 10), like ~1.92 Ga now obtained for the Early Svecofen- that of the VMB, has also been obtained from nian (DE ) deformation therefore reinforces the this volcanic belt. suggestion (Rutland et al., 2001, p. 256) that it Thus, the OMB and the pre-1.96 Ga do- could be related to the collision that produced mains adjacent to it seem to offer potential the Lapland Granulite Belt to the north of the source areas for 2.10–1.98 Ga (and Archaean) Karelian Province (Fig. 1). zircons in sediments of the younger stage of Svi- The dating of VD1 also strongly supports onian Basin development, as exemplified yb the the marginal basin accretion hypothesis for the VMB. However, the indications are that simi- VMB, and rules out various earlier tectonic in- lar, but undated, rocks extend northwards to the terpretations of the VMB based on the assump- Gulf of Finland (Fig. 1; Claesson et al., 2001, tions that the protolith sediments and the sub- Fig. 1; Bogdanova et al., 1996, Fig. 7.2a). Ko- sequent deformation and metamorphism were stinen et al. (1996, pp. 51–52) consider that the post-1.91 Ga. The sediments in the north of the structures south of the Gulf appear to represent VMB can no longer be interpreted as a fore arc a continuation of Svecofennian structures from (Lahtinen 1996, p. 43), or accretionary complex the southern Finland coastal area. They regard (Kähkönen, 1999, pp. 26–27) closely related in the land–Paldiski–Pskovshear zone, which time to the TSB; and the main structures in the forms the SW boundary of this pre-1.96 Ga do- VMB, including the suggested sutureinthe main (and the NE boundary of the younger Be- south of the VMB (e.g. Kähkönen et al., 1994, larus-Baltic Granulite belt, BBG, Fig. 1) where Fig. 1; Lahtinen, 1996, p. 95), can no longer be it crosses the Gulf of Finland, as one of the most interpreted as due to accretion at ~1.89 Ga. important internal boundaries within the Sve- The inference that a ~1.89 Ga suture might cofennian Domain. The available geological ev- exist between the TSBand the Hämeenlin- idence therefore suggests that a similar older Pal- na-SomeroVolcanic Belt (HSB, Hakkarainen, aeoproterozoic domain may form the proposed 1994) to the south (Fig. 15) was based on the early stage of Svionian Basin development in the assumed age of the VMB, and on the presence Svecofennian province proper (e.g. beneath the of MORB-affinity volcanics and DM-affinity LGSM and the CFGC, Fig. 15) and thus offer a mafic plutonics near its sheared southern mar- more local source for the VMB detrital zircons. gin. We have now shown that the VMB is older than ~1.92 Ga and that such maficro cks may be as old as 1.97 Ga (section 7.2.). If they are 8.2.The Early Svecofennian deformation indicativeofasuturerather than within-plate The most significant of our conclusions is that extension, subduction may have occurred dur- the metamorphic event in the VMB at ~1.92 ing closure of the basin at ~1.92 Ga, but if so, it Ga post-dated deposition of the sedimentary protoliths of the VMB and correlates with the observed VD metamorphism (section 7.2). In Fig. 15. Geological mapofsouthern Finland (simplified 1 from Korsman et al., 1997) showing the principal volcan- the Skellefte district, D 1 could be dated only ap- ic sequences and other tectonic elementsdiscussed in proximately by the available evidence as old- the text, and also showing the location of Fig. 3. Pre-1.91 Ga deformation and metamorphism in the Palaeoproterozoic Vammala Migmatite Belt… 131 132 Rutland,R.W.R.,Williams,I.S. and Korsman,K. occurred before the deposition of the overlying (e.g., Ekdahl, 1993; Lahtinen, 1994, Nironen, TSB and HSB. Thus, there now seems to be lit- 1997). It could also be aconsequence of clo- tle justification for postulating subduction and sure of the Svionian marginal basin under the an arc-accretionaryepisode at 1.89–1.88 Ga marginal basin accretion hypothesis, since this (e.g. Lahtinen, 1994, p. 109 & Fig. 29), or later might also involvesubduction at the inboard (Väisänen et al., 2002). margin. Direct evidence for the orientation of It is evident that the VMB is largely responsi- the active margin presumed to be outboard of ble for the crustal scale Southern Finland electri- the marginal basin to the west, or for the orien- cal conductive belt (e.g. Korja et al., 1993, Figs. tation of subduction related to it, is tenuous. D E 5 & 8). This and other such belts were previous- fold trends usually have a substantial meridion- ly interpreted as resulting from collisions post- al component but are rather variable and affect- 1.9 Ga. As in the Skellefte area (cf. Rutland et ed by later deformation. al.2001a, pp.233–234), these belts can now APhanerozoic analogue for the early evolu- be related to the pre-1.90–1.92 Ga metamor- tion of the Svionian basin may be provided by phic complexes. The continuity of these belts the Japan Sea, which was formed rapidly be- provides good evidence of the continuity of the tween 23 and 15 Ma (e.g. Okamura et al., 1998), metamorphic complexes (Fig. 2), and there- is some 900 km wide, and is only partly floored fore,for the concept of an extensivepre-1.92 by oceanic crust. The Lachlan belt in Australia Ga Svionian Basin. Nironen at al. (2002) do is an analogue for the basin filland for the sub- not show a boundary between their Pirkanmaa sequent tectonic evolution of the Svionian Ba- (including the VMB) and Pohjanmaa (WBB in sin. Ordovician turbidites extend right across the Fig. 2) belts. Koistinen et al. (1996, p. 28 and Lachlan belt and, with the exception of the east- Fig. 1) also suggest that extensions of their To- ernmost zone, represent the fillofamarginal ba- nalite Migmatite Zone (of which the VMB is a sin or basins that were originally morethan 1000 part) link eastwards to the Haukivesi-Savonlin- km wide (e.g. Fergusson, 2003). na area and to the Sortavala region northwest of Thesubsequent historyofthe Lachlan belt Lake Ladoga. was different in different areas and thereisev- We therefore now infer that the main crus- idence for diachronous deformation. Howev- tal thickening took place at ~1.92 Ga during er,the most significant deformation of the se- the closure of the younger stage Svionian Basin, quence (traditionally known as the Benambran rather than at ~1.89 Ga as proposed in the arc- Orogeny, cf. Gray et al., 1997; VandenBurg, accretionary model. Subduction zones may have 1999; Collins and Hobbs, 2001; Willman et al., been involved in the accretion process and pre- 2002), which closed the marginal basin, took

1.91 Ga sutures and/or major D E thrust struc- place in the late Ordovician to early Silurian. It tures may therefore be present. Rutland et al., is comparable to the ~1.92–1.90 Ga (DE )defor- (2001a) suggested that the earliest deformation mation and metamorphism that closed the Svi- phase (DE1) could perhaps be directly responsi- onian Basin in the Svecofennian province. The ble for the deep seismic reflectors,interpr eted as style of the early deformation in the crust of the a fossil subduction zone, in BABEL profiles 3 Lachlan belt, including major lowangle thrusts, and 4 (BABEL Working Group, 1990; 1993). has recently been revealed in the NE of the belt The NE dip of this zone is consistent with the by seismic reflection studies (Glen et al., 2002). direction of D E1 thrusting onto the adjacent Butthe natureofthe basement and the specif- Karelian craton. In the arc accretion models this ic role of subduction in producing the tectonic earliest deformation has been attributed to ac- pattern haveyet to be resolved (e.g. Gray et al., cretion of the earliest arc complex through sub- 2002; Cayley et al., 2002). This emphasises that duction either towards or away from the craton much further workisrequired to establish the Pre-1.91 Ga deformation and metamorphism in the Palaeoproterozoic Vammala Migmatite Belt… 133 histories and mutual relations of the various pro- in the early deformation history. The intrusion posed Svionian metamorphic complexes, before of the potassium-rich granites of the LSGM in plausible models of pre-1.91 Ga Svecofennian the Turku district is attributed to the local D 3 . evolution can be presented. Akey related ques- The earlierD1 /D2 deformations (op. cit.) in the tion is the extent to which the main plutonic do- gneisses appear to have taken place under sim- mains adjacent to the metamorphic complexes ilar metamorphic conditions and to have sim- contain evidence of apre-1.91 Ga history. ilar relationships to the early granitoids (here ~1.89–1.87Ga), as havebeen described from the VMB. There is therefore a possibility that 8.3. Evidence for pre–1.91 Ga structures in the the earliest bedding parallel biotite + sillimanite CFGC and LSGM foliation (mostly obliterated by younger events, Various areas of veined gneisses, similar to op. cit., p. 183) is a pre-1.91 Ga structure like those in the VMB, havebeen mappednear that in the VMB. to the southern and eastern boundaries of the It has been noted that the volcanic belts of CFGC. In the SE corner of the CFGC, Niro- SW Finland show a general transition from bi- nen (1995) reports an age of 1906±10 Ma for modal sequences in Enklinge-Orijärvi, through a medium grained tonalite, which he considers calc-alkaline associations, to morealkaline post-dates deformation and metamorphism in volcanics in Tampere(Ehlers &Lindroos, older tonalite/granodiorite and in paragneisses. 1997). Broadly the thick basaltic and andesit- This age is now considered to be false because ic sequences accompanied by thick turbiditic se- the conventional data are heterogeneous, prob- quences in the TSB and HSB form a northern ably due to inherited zircon (H. Huhma, pers. domain. This domain contrasts with a 100 km- comm., 2004). Nevertheless, the geological re- wide southern domain (Fig. 15), where a thin lationships indicate that the older deformation succession contains pillow lavas with EMORB and metamorphism might be the same as in the characteristics associated with shallowmarine VMB and further work is desirable. sediments including marbles and i ronforma- The late Svecofennian granite-migmatite tions. Ehlers & Lindroos (op. cit.) infer exten- zone (LSGM) of Ehlers et al. (1993; see Fig. 15) sion and rifting for this southern domain, and it is a zone of high-grade metamorphism, migma- presumably occurred in a previously more stable tites and S-type granites dated between 1.84 and tectonic environment than that of the northern 1.80 Ga. Older tectonic relationships are there- domain. The nature of the basement in which fore difficultto decipher. High-grade gneisses in this rifting occurred has not been established. the Tu rku district (about 50 km NW of Kemiö The area of bimodal volcanics south and township (Fig. 15) are associated with a southwest of Hyvinkää (Fig. 15) has been called the westerlyextension of the Hämeenlinna-Some- Kemiö-Mäntsälä Belt (KMB, e.g. Kähkönen et ro volcanic belt (HSB). It has been assumed that al., 1994) and is generally regarded as an east- all the sequence was involved in all the defor- ward continuation of the Bergslagen ore prov- mation history so that a date of 1888±11 Ma ince in Sweden, so that the felsic volcanics (Vaasjoki, 1994) from the upper partofthe would be younger than ~1.9 Ga (e.g. Lindroos HSB has been considered to provide not only & Ehlers, 1994, Fig. 1; Lundström et al., 1998). a minimum age for the sequence (Väisänen and This has been confirmed in the Orijarvi area, Hölttä, 1999, p. 182) but also a maximum age where IDTIMS U-Pb zircon analyses from the for the onset of deformation (op. cit., p. 205). bottom and top of the felsic succession yielded However, the close similarity of the HSB evo- ages of 1895.3±2.4 Ma and 1878.2±3.4 Ma re- lution to that of the TSB (see section 8.4 be- spectively (Väisänen & Mänttäri, 2002). How- low), suggests that the HSB was not involved ever, only upright folds affect these rocks, and 134 Rutland,R.W.R.,Williams,I.S. and Korsman,K. primary volcanic and sedimentary features are rifting at ~1.95– 1.96 Ga may have split the preserved (Colley & Vestra, 1987, p. 96). Thus older stage Svionian complex so that the the felsic sequence is similar in age and structur- younger Svionian Basin beneath the VMB al style to the TSB, and it seems possible that, as may havebeen floored by newoceanic crust, in the Tampere district, earlier deformation took or by older crust of much reduced thickness. place in underlying rock groups before the fel- This might help explain the absence of younger sic volcanics were erupted. The observation that granites in the VMB, while they are prominent polymict breccias contain blocks of gneiss, rein the older crustal areas to north and south. garded by Colley and Vestra (op. cit. p. 97) as possible Archaean basement, but which might also be pre-1.91 Ga or pre-1.98 Ga Palaeoprot- 8.4.The post- 1.91 Ga volcanic sequences erozoic gneisses, offers an avenue for future re- The volcanic sequence of the TSB indicates an search. extensional period between the early deforma-

Furthermore, the stratigraphy of the KMB tion of the VMB (VD 1 ) and the later deforma- is controversial, and it has been suggested that tion of the TSB (TD 1 ) and this is well bracket- the thin marbles with associated iron formations ed between ~1.92 Ga and ~1.88 Ga (cf. Rutland and cherts, referred to above, may form a mark- et al., 2001a, b) for the Skellefte district (SD) er horizon, separating two distinct rock groups in Sweden. Thereisthereforea40m.y.peri- (Lindroos, 1990, reported by Kähkonen et al., od, during which the TSB sequence was depos- 1994). Thus the mafic rocks indicative of rift- ited, before the ~1.88 Ga folding and metamor- ing (Ehlers et al., 1986) in the Nagu-Korpo area, phism. There is no evidence of local subduction about 40 km WNW of Kemiö township, (Fig. during this period, although subduction pre- 15), may be significantly older than the dated sumably continued beneath an active margin felsic volcanics. They resemble recent within- outboard of the closed Svionian marginal basin. plate lavas (cf. Haveri) and are closely associated The analogous Late Silurian and Early Dev- with the Turku gneisses discussed above. Their onian historyinthe Lachlan belt in Australia possible correlation with the maficvo lcanism in is one of crustal extension producing shelves, the VMB needs to be explored. troughs and inferred half grabens in the under- We surmise therefore, on the basis of these lying deformed Ordovician complex (e.g. Glen indications of pre-1.91 Ga structures and rock et al., 2002; Willman et al., 2002) well inboard units, that the younger stage Svionian Basin of the contemporaneous active margin. This ex- may well haveextended through the areas of tensional episode of about 50 m.y. corresponds the CFGC and LSGM. We have also suggested to the main period of granitoid emplacement above (section 8.1) that both areas may be and volcanism in the belt (see e.g. Zen, 1995; underlain by the older stage Svionian basin Collins, 1998; Chappell et al., 2000). Thus, it accreted before ~1.96 Ga. The accretion of the occupied a similar time span to the ~1.92–1.88 older and younger stage Svionian basins there- Ga extensional period of granitoid emplace- fore provides a possible explanation for the old- ment and volcanism in the TSB. er crust postulated on geochemical grounds in The TSB and other volcanic sequences there- both of these domains (e.g. Lahtinen & Huh- foreneed to be re-evaluated as the preserved ma, 1997; Rämöet al., 2001). The contrast, remnants of awidespread episode of exten- noted above,between the northern and sional volcanism, which followed the accretion southern domains during the post-accretionary of the Svionian marginal basin. The close sim- volcanism, also suggests that the VMB differs ilarities in age and in qualitative evolution be- in crustal character from the LSGM and the tween the Skellefte (SD), Tampere (TSB) and CFGC. We thereforefurther speculate that Hämeenlinna-Somero(HSB) volcanic belts Pre-1.91 Ga deformation and metamorphism in the Palaeoproterozoic Vammala Migmatite Belt… 135

(Fig. 15) appear to favour the post-accretion- groups, where the presence of a basement do- aryinterpretation. In all three areas the lower main preserving N-S trends is indicated (e.g. parts of the sequences are characterized by fel- Rutland et al., 2001b). South-side up shears sic volcanics closely associated with greywack- separate the SD from the pre-1.9 Ga migma- es and pelites and are followed by tectonic ac- tites of the Robertsfors Group to the south and a tivity leading to the deposition of conglomer- similar shear separates the TSB from the VMB. ates containing clasts derived from the under- Immediately north of the TSB, the E-W defor- lying sequence or from granitoids of similar mation also affects the adjacent CFGC, where age. The lower felsic volcanics in the TSB, with volcanics and sediments, presumably correla- published ages of 1904±4 Ma, 1898±4 Ma and tive with the TSB, are also preserved in the Hir- 1892±3 Ma (Kähkönen et al., 1989; 2004) ap- silä Belt (Fig. 15). Further north, similar volcan- pear to be slightly older than the published ages ics again appear in the Haukkamaa area (Tiain- of 1888.8±4.2 Ma and 1882±8 Ma in the SD en and Kähkönen, 1994), but NW trends dom- (Billström and Weihed, 1996). In the Skellefte inate in remnant supracrustal belts and in the district, the change in depositional character has early granitoids in the intervening area, and been attributed to uplift in response to relaxa- may again be indicative of a less deformed base- tion of crustal extension and the emplacement ment domain. One of several ~W-trending belts of large intrusions to shallow crustal levels (Al- of gabbroic intrusions locally marks the south- len et al., 1996) and it is likely that this explana- ern boundary of this domain (see Kuru 1:100 tion has more general application. 000 map sheet 2213, Matisto, 1960), suggest- In all three areas this tectonic activity is also ing a relationship to extensional structures initi- followed by more maficupper volcanic sequenc- ated during the deposition of the TSB. es, probably reflectingre newed extension. In the The VMB can be regarded as a major TD1 TSB these mafic rocks are limited to small are- anticlinorium separating the synclinal structures as in the axial depressions of the main syncli- containing the similar stratigraphic sequences of nal hinge zone, and one date of 1889±5 Ma the TSB to the north and the HSB to the south has been obtained (Kähkönen et al., 1989). (Fig. 15). In both cases, the main occurrence of In the HSB a meta-andesite has been dated at turbiditic sediments associated with the volcan- 1888±11 Ma (Vaasjoki, 1994). In the SD sig- ic sequences occurs adjacent to the VMB. In the nificantlyyo unger dates of 1878±2 Ma, 1876±3 HSB, Hakkarainen (1994) considered that the Ma and 1874.9±3.7 Ma havebeen obtained extensive Häme Group was erupted from an E- (Skiöld et al., 1993). The available evidence in- W fissure system. Thus, it seems likely that in dicates that the folding in the SD (~1.86 Ga, the TSB and HSB, as in the SD (Rutland et al., Rutland et al., 2001b) was also somewhat lat- 2001a, p. 230), the post-TSB fold structures er than in the TSB (1.88 Ga, Nironen, 1989b). areparalleling earlier W-trending extensional The age of the folding is not established in the or transtensional rift structures in the pre-1.91 HSB but it is notable that the granitoids dated Ga basement. in the western areas of the HSB at ~1.87 Ga are The necessaryre-appraisal of the petrolo- about 10 Ma younger than related granitoids in gy and geochemistry of the volcanic sequences the CFGC and TSB (Väisänen et al., 2002). is beyond the scope of this paper, but the pro- Major upright synclines with E-W trend longed pre-1.9 Ga crustal history now proposed characterise all three areas, and the volcanic se- and the crustal thickening prior to the volcan- quences were probably originally much more ism appear to fitwe ll with the evolved character extensive. In the SD the younger maficsequenc- of the volcanics and the evidence of a pre-1.9 Ga es are extensive to the north of the main syn- component in both the volcanic and plutonic clinal structure, in the Vargfors and Arvidsjaur 1.90–1.87 Ga suites outlined in section 1 above. 136 Rutland,R.W.R.,Williams,I.S. and Korsman,K.

Characteristics transitional towards within-plate and LSGM might also be explained as thicker basalts are also sometimes present (Kähkönen, areas of the accreted older (~2.1–1.98 Ga) stage 1987, 1994; Viluksela, 1994). Possibly, arc char- of Svionian Basin evolution. The younger grey- acteristics could have been inherited from slabs wackes in the post-~1.91 Ga Bothnian Basin are subducted during the ~1.92 Ga closure of the largely the consequence of erosion of the pre- Svionian Basin. Certainly,the facies variation 1.92 Ga complexes. within the belts has no necessary relation to the The wide regional distribution of 1.90–1.87 W-trending extensional or fold structures, or to Ga plutonic and volcanic rocks (extending into previously postulated subduction zones. For ex- the Archaean basement of the Karelian Prov- ample, a N-trending zonation has been report- ince), the exclusively HT/LP metamorphism ed in the volcanic province north of the Skel- and the extensivemigmatisation all indicate a lefte district (Perdahl & Fritsch, 1993) and also very high heat input over a wide area between from the Bergslagen district (Lagerblad, 1988). 1.90 and 1.87 Ga, which is plausibly explained We also note that some examples of calc-alka- by the ~40 m.y. extensional episode following line magmatism elsewhere are not as closely re- the main episode of crustal thickening. The un- lated to subduction zones in time or space as is usually thick and high-velocity lowercrust in often assumed (e.g. Chappell et al., 2000; Col- some areas (Korja et al., 1993), and notably be- lins, 2002; Bryan et al., 2003). neath the CFGC, may also be largely the consequence of underplating during this extensional episode (cf. Etheridge et al., 1987; Rutland 8.5. Concluding Comment et al., 2001b). However, both the Svionian and The results discussed here, in conjunction with younger extensional events may also have con- earlier work in the Skellefte region, show that tributed to the underplating (e.g.at~1.8 Ga, the hypothesis of rapid arc-accretion post-1.9 Peltonen & Mänttäri, 2001). Ga for the greater partofthe Svecofennian The Palaeoproterozoic north Australian oro- Province can no longer be sustained. The availa- genic province and the Palaeozoic Lachlan belt ble evidence indicates a long earlier history with in eastern Australia arealso characterised by two main episodes of accretion between periods thick high-velocity lower crust and by HT/LP of basin formation at ~2.1–1.98 Ga (older Svi- metamorphism and abundant granitoids (e.g. onian), ~1.97–1.92 Ga (younger Svionian) and Rutland, 1982, 1997). The strong similarities ~1.91–1.86 Ga (Bothnian). We suggest that the in tectonic evolution between the Svecofennian alternative hypothesis of marginal basin accre- Province and the Lachlan belt suggest that they tion discussed above provides amoresatisfa- developed in similar tectonic settings. tory explanation of the main features of the Sve- cofennian orogenic province. The occurrence Acknowledgments of older migmatite complexes beneath the vol- We thank Hannu Huhma and Mikko Nironen for canic belts is explained in terms of the young- constructive reviews that resulted in major improve- er stage Svionian marginal Basin (~1.97–1.92 ments in the manuscript. Williams and Rutland would Ga), closed during the the major ~1.92 Ga tec- like to acknowledge their debt to numerous Finnish tono-thermal event, which was responsible for geologists, and especially to co-author Korsman, who have helped them develop some, albeit imperfect, un- the main episode of crustal thickening, proba- derstanding of Finnish Precambrian geology,both bly involving pre-~1.91 Ga subduction zones. in the field and through the exceptionally compre- The pattern of this Early Svecofennian orogenic hensive body of literature that is available. Particular phase remains to be explored, and exotic micro- thanks are due to Timo Kilpeläinen whose work provided the essential basis for this study and who has as- continents may have been accreted at that time. sisted throughout; and to, Carl Ehlers, Gabor Gaál, However, the older crustal areas of the CFGC Asko Kontinen, Tapio Koistinen, Jukka Kousa, Raimo Pre-1.91 Ga deformation and metamorphism in the Palaeoproterozoic Vammala Migmatite Belt… 137

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