U–Pb and Sm–Nd isotopic constraints on the evolution of the Paleoproterozoic Peräpohja Belt, northern Finland

Precambrian Research 266 (2015) 246–259

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U–Pb and Sm–Nd isotopic constraints on the evolution of the Paleoproterozoic Peräpohja Belt, northern Finland Jukka-Pekka Ranta a,∗ , Laura S. Lauri b , Eero Hanski a , Hannu Huhma c , Yann Lahaye c , Erkki Vanhanen d a

Oulu Mining School, University of Oulu, P.O. Box 3000, 90014, Finland Geological Survey of Finland, P.O. Box 77, 96101 Rovaniemi, Finland c Geological Survey of Finland, P.O. Box 96, 02151 Espoo, Finland d Mawson Oy, Lantontie 34, 95680 Lohijärvi, Finland b

a r t i c l e

i n f o

Article history: Received 10 December 2014 Received in revised form 16 May 2015 Accepted 18 May 2015 Available online 23 May 2015 Keywords: Geochronology Paleoproterozoic Jatuli Kaleva Peräpohja Belt Finland

a b s t r a c t Ten localities, representing metasedimentary and igneous rocks from the northern part of the Paleoproterozoic Peräpohja Belt, Fennoscandian Shield, were sampled for zircon U–Pb dating by LA-MC-ICP-MS and Sm–Nd analysis by TIMS. The results revealed two phases of granitic magmatism that took place at 1.99 Ga and 1.79–1.77 Ga. The former age (1989 ± 6 Ma) was given by the Kierovaara granite pluton representing the first example of felsic plutonism of this age in northern Finland, though A-type felsic volcanic rocks of the same age have previously been recognized. The strongly negative initial εnd values obtained for granitoids, ranging from −3.9 to −11.4, are compatible with the previous Sm–Nd data, indicating the presence of Archean basement in northern Finland at the time of granite emplacement. A garnet-whole rock Sm–Nd age of 1752 ± 14 Ma was measured for a mica gneiss, constraining the time of metamorphic cooling down to the closure temperature of garnet. Detrital zircon grains from quartzite samples belonging to the Jatulian Palokivalo Formation and a stratigraphically unassigned lithodemic unit, the Mellajoki Suite, yielded solely Archean zircon populations. In contrast, our sample from the Martimo Formation representing Kalevian sedimentation revealed a 66% Paleoproterozoic detrital zircon component with the youngest zircon grains showing an age of ca. 1.91 Ga. The new age data allow us to constrain the depositional time of this sedimentary unit between ca. 1.91 and 1.88 Ga, the latter being the time of the synorogenic plutonism. Our geochronological data coupled with recent data from other Karelian belts suggest that the depositional time period for rocks assigned to the Kaleva spans ca. 150 Ma and the simple division of these rocks into two non-coeval units, the Lower and Upper Kaleva, has become problematic. © 2015 Elsevier B.V. All rights reserved.

1. Introduction The Peräpohja Belt (PB) is one of the major Paleoproterozoic supracrustal rock sequences in the Fennoscandian Shield, representing part of the Karelian volcano-sedimentary cover deposited on the Archean cratonic basement (Fig. 1). It records a typical evolution of a Karelian depositional basin starting with clastic, relatively mature sedimentary rocks and subaerial mafic volcanic rocks deposited on a rifted Archean cratonic basement. Gradual deepening of the basin was accompanied by accumulation of carbonate rocks displaying a heavy carbon isotope excursion diagnostic of Jatulian carbonate rocks (Karhu, 1993). The upper part of

∗ Corresponding author. Tel.: +358 440557857. E-mail address: jukka-pekka.ranta@oulu.fi (J.-P. Ranta). http://dx.doi.org/10.1016/j.precamres.2015.05.018 0301-9268/© 2015 Elsevier B.V. All rights reserved.

the sequence is dominated by deep-water turbiditic mica schists and black shales. This general evolution, which spans several 100s of million years, is similar to what has been observed in other Karelian belts in the eastern part of the Fennoscandian Shield (i.e. Laajoki, 2005; Hanski and Melezhik, 2012). Due to this similarity, shield-wide correlations have often been made using traditional stratigraphic names such as Sumi, Sariola, Jatuli, Ludicovi, and Kaleva (e.g., Meriläinen, 1980; Ojakangas et al., 2001). Recently, these names were redefined as systems with a geochronological connotation (Hanski and Melezhik, 2012). In the Peräpohja Belt, the above-described supracrustal evolution constitutes Sariolian, Jatulian, Ludicovian and Kalevian rocks. The evolution of the Karelian supracrustal belts commenced with the Sumian System (2505–2430 Ma), comprising volcanogenic rocks ranging from komatiites to rhyolites in composition with local polymictic volcanogenic breccias and other volcanoclastic

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Fig. 1. Karelian supracrustal belts in eastern and central Finland (modified after Laajoki, 2005). The geological map of the framed part of the Peräpohja belt is shown Fig. 2.

rocks (Hanski and Melezhik, 2012). The emplacement of ca. 2505–2430 Ma layered gabbronorite intrusions is assigned to the Sumian system, while the Sariolian system (2430–2300 Ma) rocks postdate these intrusion. The latter are represented by andesites to komatiitic volcanic rocks and immature sedimentary rocks (i.e. polymictic conglomerates) related to the Huronian-age glaciation, varying in thickness up to 600 m and resting either on the igneous rocks of the Sumian rocks or on Archean basement rocks (Hanski and Melezhik, 2012). The Jatulian system (2300–2060 Ma) with a thickness up to 2000 m is composed of mafic volcanic rocks, mature arenites, quartzites, conglomerates and stromatolitic carbonate rocks with a Lomagundi-Jatuli enriched 13 C-isotopic signature (Karhu, 1993; Melezhik et al., 2013). The Ludicovian system (2060–1960 Ma) lies mostly on Jatulian rocks and represents mafic volcanic rocks, accumulations of organic carbon-rich metasediments, the Earth’s earliest known phosphorites and first formation of 13 C-depleted diagenetic carbonate concretions (Hanski and Melezhik, 2012). The Kalevian system (1960–1900 Ma) contains turbiditic qreywackes and black shales deposited on the south-western margin of Karelian craton, with the estimated thickness varying between 3000 and 5000 m (Hanski and Melezhik, 2012). Kontinen (1987) defined

the Lower Kalevian rocks as autochthonous–parautochthonous rock units whereas the Upper Kaleva is dominantly allochthonous and encloses fragments of ophiolite complexes (see also Lahtinen et al., 2010). Fig. 1 shows the distribution of Karelian successions in eastern and central Finland, as defined by Laajoki (2005). In this simplified map, the Jatulian rocks include Ludicovian rocks that in Finland are equivalent to rocks that are traditionally called “marine Jatulian” (Meriläinen, 1980; Hanski and Melezhik, 2012). Recent studies in the Paleoproterozoic North Karelia Belt in eastern Finland have shed new light on the lithological characteristics, provenance and age of deposition of the two Kalevian units. Lahtinen et al. (2010) presented geochronological data indicating that detrital zircon grains in the Lower Kaleva were derived almost solely from Neoarchean rocks of the Karelian craton whereas the Upper Kaleva contains a large Paleoproterozoic zircon population from sources as young as 1.91–1.95 Ga, with the Lapland-Kola Orogen considered one potential source of these grains. However, this simple picture, which is reasonable given the autochthonous versus allochthonous nature of the two Kalevian units, has become more complicated since Lahtinen et al. (2013) published additional isotopic data from the North Karelia Belt, revealing the presence of

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ca. 1.91–1.92 Ga zircon grains in rocks units that have earlier been assigned to the Lower Kaleva. The lithostratigraphy of the Peräpohja Belt is rather well-known for the most part of the belt (Perttunen, 1985, 1989; Perttunen and Hanski, 2003; Kyläkoski et al., 2012). However, the exact depositional age of the upper part of the supracrustal succession consisting of Kalevian turbiditic metasediments and black shales has remained elusive due to the difficulties in determining direct radiometric ages for these sedimentary rocks. It has only been possible to state for certain that the pelitic rocks of the belt are younger than the underlying c. 2.1 Ga metabasalts and older than the crosscutting, ca. 1.88 Ga synorogenic monzonitic plutons (Perttunen and Vaasjoki, 2001). The correlation of the Peräpohja mica schists and black shales with the Lower and Upper Kalevian units elsewhere in Finland has also been unclear though the Lower Kaleva connection is normally assumed mainly because no definite unconformities between the Kalevian and Jatulian rocks have been observed (Laajoki, 2005). The stratigraphy of the upper (northern) part of the Peräpohja Belt is further complicated by the recognition of lithological units at the northern margin of the belt that do not correspond to conventional Karelian supracrustal rocks, notably 1.98 Ga felsic tuffs with an A-type geochemical signature (Hanski et al., 2005). Moreover, the metamorphic grade and degree of deformation increases toward the northern margin of the belt, leading eventually to transition of the supracrustal belt, over a short distance, into the migmatitic Central Lapland Granitoid Complex. These features led Laajoki (2005) to separate the northern margin from the rest of the Peräpohja Belt in his map (see Fig. 1). Impetus to this study came from the recent discovery of an extensive gold-mineralized zone in the northwestern part of the Peräpohja Belt, encompassing the Rompas Au-U and Rajapalot Au prospects (Vanhanen et al., 2015). Due to polyphase deformation, amphibolite facies metamorphism, a high degree of alteration and a limited number of outcrops in this part of the belt, the primary character of the rock units and their stratigraphic relations are relatively poorly known. To improve our understanding of the geological development of the gold-bearing area and the whole Peräpohja Belt, we collected samples from Jatulian and Kalevian sedimentary rocks and a lithodemic unit close to the mineralized area for determining the U–Pb age spectrum of their detrital zircon grains and measuring their whole-rock Sm–Nd isotopic compositions. We also sampled granitic intrusions and leucosomes of migmatites within the Peräpohja Belt and its northern margin in order to obtain new age constraints for the belt and allow comparison of the granitic magmatism in the Peräpohja Belt and the Central Lapland Granitoid Complex. Our new isotopic data confirmed the correlation between quartzite units in different parts of the belt, provided new evidence for the presence of Upper Kalevian pelitic sedimentary rocks in the belt, and led to the discovery of a previously unrecognized granitic magmatic phase in northern Finland.

2. Regional geology The Peräpohja Belt (PB) comprises a sequence of Paleoproterozoic supracrustal rocks deposited unconformably on the Archean Pudasjärvi Complex. The basement complex is exposed on the southern side of the PB (Fig. 2) and consists of Archean granitoids, remnants of greenstone belts and Paleoproterozoic mafic layered intrusions. In the north, the PB is bordered by the Paleoproterozoic Central Lapland Granitoid Complex (CLGC). The western margin of the PB is defined by a N-S-trending shear zone running close to the Tornionjoki River valley. The maximum depositional age for the PB is constrained by the 2.44 Ga mafic layered intrusions, which were partly eroded when the oldest sediments started to accumulate

(see Iljina and Hanski, 2005 and references therein). On the other hand, a minimum age is determined by the ca. 1.88 Ga synorogenic monzonite intrusions of the Haaparanta (Haparanda) Suite, which cut the youngest metasediments of the PB (Perttunen and Vaasjoki, 2001). Perttunen et al. (1995) divided the supracrustal rocks of the PB into two major lithostratigraphic units, the Kivalo and Paakkola Groups, which are correlative, respectively, to the Jatulian and Kalevian rocks, and subdivided them into eleven formations. Hanski (2001) and Perttunen and Hanski (2003) presented a more detailed subdivision of the Kivalo Group. Later, two previously unrecognized rock units were defined as new formations in the lithostratigraphy of the PB (Hanski et al., 2005; Kyläkoski et al., 2012). The current interpretation of the stratigraphy is shown in Table 1. The lower part of the Kivalo Group consists mainly of quartzites, mafic volcanic rocks and dolomites. The basal conglomerate of the Sompujärvi Formation forms the base for the PB. It is overlain by amygdaloidal mafic lavas of the Runkaus Formation, which has a minimum age of ca. 2250 Ma (secondary titanite Pb–Pb age; Huhma et al., 1990). Quartzites of the Palokivalo Formation were deposited on the Runkaus Formation. They are cut by ca. 2.22 Ga mafic differentiated sills (Perttunen and Vaasjoki, 2001; Hanski et al., 2010), which define a minimum age for the quartzites. Published multigrain zircon TIMS data indicate exclusively or dominantly Archean sources for the zircon grains in the Palokivalo Formation (Hanski et al., 2001; Perttunen and Vaasjoki, 2001). The Kaisavaara Formation is composed of sericitic quartzites. The stratigraphic position is still problematic, but Perttunen and Hanski (2003) correlated it with the Palokivalo Formation. One of our samples comes from the Mellajoki Suite (Fig. 2), a lithodemic unit in the N-W margin of the PB, comprising quartzites, mica schists and mica gneisses. The goldmineralized area is located just east of these quartzites. So far, the correlation of the highly deformed Mellajoki Suite with quartzites of the Kivalo Group has been uncertain and therefore the former has been classified as a lithodemic unit (Perttunen and Hanski, 2003). The quartzites are overlain by the Petäjäskoski Formation. It consists mainly of phlogopite-sericite schists and albitic schists, which are considered to have originated in an evaporitic depositional environment (Kyläkoski et al., 2012). The Petäjäskoski Formation is overlain by the continental flood basalts of the Jouttiaapa Formation showing, for continental flood basalts, abnormal, strongly LREE-depleted chondrite normalized REE patterns (Perttunen and Hanski, 2003; Hanski, 2012). Whole-rock Sm–Nd data from the Jouttiaapa basalts suggest an age of 2105 ± 50 Ma (Huhma et al., 1990; Hölttä et al., 2003), which is consistent with the U–Pb zircon age of 2140 ± 11 Ma obtained for a potentially comagmatic mafic sill (Kyläkoski et al., 2012). The upper part of the Kivalo Group consists of quartzites of the Kvartsimaa Formation and intervening dolomites and mafic tuffites belonging to the Tikanmaa, Poikkimaa, Hirsimaa, Rantamaa, and Lamulehto Formations. A zircon U–Pb age of 2106 ± 8 Ma has been published for a mafic tuffite from the Hirsimaa Formation (Karhu et al., 2007). The Paakkola Group comprises mica schists, black schists, pillow basalts, and mafic and felsic tuffs. In terms of chronostratigraphy, the lowermost unit is the Väystäjä Formation (Table 1), although it is not observed in the field at the contact of the Kivalo and Paakkola Groups, as it occurs in a separate area between the Martimo Formation and the granites of the Central Lapland Granitoid Complex (Fig. 2). It is composed mostly of pillowed basalts, but dolomites and felsic porphyries also occur within the unit. Perttunen and Vaasjoki (2001) reported a U–Pb age of 2050 ± 8 Ma for a felsic porphyry unit of the Väystäjä Formation. The Väystäjä Formation is spatially associated with Martimo Formation metasediments (Fig. 2), but the contact between these two formations is probably tectonic. The felsic and mafic tuffs of the Korkiavaara Formation are among the youngest supracrustal rocks of the PB as shown by the

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Fig. 2. Geological map of the western part of the Peräpohja Belt (modified after Bedrock of Finland – DigiKP, Digital map Database, version 1.0, Geological Survey of Finland, 2011, http://www.geo.fi/en/bedrock.html and Kyläkoski et al., 2012). For location of the map, see Fig. 1. Black stars represent sampling sites.

zircon U–Pb age of 1973 ± 11 Ma measured for felsic tuffs (Hanski et al., 2005). The Pöyliövaara Formation comprises mica schists and black schists, which are deposited on to the Korkiavaara formation (Hanski et al., 2005). Uppermost in the stratigraphy is the Martimo Formation, which based on the detrital zircon populations measured in this study represents the youngest supracrustal rocks of the area. The Martimo Formation is composed of metagraywackes, mica schists and black schists that frequently display turbiditic structures (Perttunen, 1989). The rocks of the PB have undergone polyphase deformation and are metamorphosed under greenschist facies conditions in the southern parts of the belt. The metamorphic grade is higher in the northern and eastern parts, where amphibolite facies conditions prevail. In the northeastern marginal zone, the rocks of the PB are locally migmatized (see Hanski et al., 2005) whereas, based on our field observations, the northern margin of the belt is most probably defined by a fault that separates the non-migmatitic rocks of the PB from the migmatitic rocks of the CLGC. 3. Samples Ten samples were collected from different units in the NW part of the PB belt to constrain the ages of granitoids and examine detrital zircon populations in sedimentary rocks. The sampling sites are shown in Fig. 2 and the main features of the sampled rocks are described below.

3.1. A2160 Männikkömaa Sample A2160 Männikkömaa is from a medium-grained, garnetbearing migmatitic mica gneiss composed of tonalitic leucosome and garnet-biotite-bearing paleosome, which are both represented in the collected sample. The mica gneiss occurs in the contact zone between the Peräpohja Belt and the Central Lapland Granitoid Complex. Its main minerals are quartz, plagioclase and biotite; apatite, chlorite, zircon, and opaque minerals are found as accessory phases. Subhedral garnets and heavily altered cordierite porphyroblasts are found mainly in the biotite-rich paleosome. Biotite is locally chloritized and plagioclase is saussuritized. Zircon is abundant especially as inclusions in biotite.

3.2. A2161 Matalavaara Sample A2161 Matalavaara was taken from a homogeneous, pink, leucocratic, medium-grained alkali-feldspar granite, which in common with sample A2160, is located in the contact zone between the Peräpohja Belt and the Central Lapland Granitoid Complex (Fig. 2). The granite intrudes migmatites represented by sample A2160, but due to poor outcrop conditions, the contacts are not exposed. Matalavaara-type granites are present throughout the northern margin of the PB from the western part of the study area to the area of the Rovaniemi town ca. 50 km eastwards, commonly forming topographic highs. Sample A2161is

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Table 1 Lithostratigraphy of the Peräpohja belt (modified after Hanski et al., 2005; Kyläkoski et al., 2012). Chronostratigraphy

Lithostratigraphy

Lithology

Age

Martimo Pöyliövaara

Mica and black schist, qraywacke Mica and black schist

<1920 Ma (U–Pb)a

Korkiavaara Väystäjä

Felsic and mafic tuff Pillowed basalt, dolomite and felsic porphyry

<1973 ± 11 Ma (U–Pb)b 2050 ± 8 Ma (U–Pb)c

Lamulehto Rantamaa Hirsimaa Poikkimaa Tikanmaa Kvartsimaa Santalampi Jouttiaapa Petäjäskoski Kaisavaara Palokivalo and mellajoki suite

Mafic tuffite Dolomite, quartzite Mafic tuffite Dolomite, phyllite Mafic tuffite Quartzite, dolomite Agglomerate, pillowed basalt Amygdaloidal basalt Mica-albite schist, dolomite, quartzite, breccias Quartzite, conglomerate Quartzite, arkosic quartzite, orthoquartzite, mica gneisses, mica schists Amygdaloidal basalt, agglomerate Conglomerate, arkose, quartzite

System

Group

Formation

Kalevian

Paakkola

Ludicovian

Jatuli

Kivalo

Sariola a b c d e f g h i

Runkaus Sompujärvi

2106 ± Ma 7 (U–Pb)d

2105 ± 50 Ma (Sm–Nd)e >2140 ± 11 Ma (U–Pb)f >2220 Ma (U–Pb)g 2250 ± 30 Ma (U–Pb)h <2440 Ma (U–Pb)i

Minimun age based on detrital zircons (this study and Lahtinen et al., 2015). Hanski et al. (2005) and Lahtinen et al. (2015). Age of a quartz porphyry (Perttunen and Vaasjoki, 2001). Karhu et al. (2007). Huhma et al. (1990) and Hölttä et al. (2007). Age of a cutting mafic sill (Kyläkoski et al., 2012). Minimum age based on cutting mafic sills (Perttunen and Vaasjoki, 2001; Hanski et al., 2010). Perttunen and Vaasjoki (2001). Maximum age based on underlying mafic layered intrusions (Huhma et al., 1990).

thorium-rich with 65 ppm Th and 5 ppm U. The main minerals are quartz, K-feldspar (microcline) and plagioclase. Biotite, apatite, zircon, allanite, uranothorite and opaque minerals are found as accessory phases. Secondary sericite is found in the fractures of the feldspar crystals and biotite has partly altered to chlorite. Hematite occurs as a dust-like pigment in the crystal boundaries and fractures. Plagioclase is locally saussuritized and the grains also show a local antiperthite texture.

are found as accessory minerals. No zircon grains were observed in thin section. Potassic feldspar and plagioclase show local perthite and antiperthite textures, and plagioclase is locally saussuritized. Sample A2164 Mustivaara S is from an outcrop consisting solely of coarse-grained tourmaline granite. It has a similar mineral composition to that of sample A2163, though garnet is not present in A2164.

3.3. A2162 Matalavaara S

3.5. A2165 Karhurommas

Sample A2162 Matalavaara S was picked from white, coarsegrained alkali-feldspar granite representing the leucosome of a migmatite in the southern part of the migmatite zone separating the CLGC and PB (Fig. 2). Among the dated granites, this granite is richest in uranium (ca. 20 ppm) and has also a high Th content (25–30 ppm). The granite has intruded into biotite-rich amphibolite with an unknown stratigraphic position. The main minerals are quartz, K-feldspar, plagioclase and biotite. K-feldspar locally forms pegmatoidal segregations in the rock. Accessory minerals are apatite, chlorite and zircon. Plagioclase is distinctly zoned and locally saussuritized and biotite is weakly chloritized.

Sample A2165 represents a garnet-bearing, psammitic mica schist from the Mellajoki Suite, located ca. 0.5 km east of the Rompas gold mineralization. The rock is fine-grained and has abundant garnet, cordierite and andalusite porhyroblasts in a groundmass consisting of quartz, biotite and chlorite. Zircon is abundant, especially as inclusions in biotite. Clusters of chlorite crystals are locally present.

3.4. A2163 Mustivaara N and A2164 Mustivaara S The two samples, A2163 and A2164, were collected from the same granitic body, a pegmatitic tourmaline granite, which seems to represent the youngest granitic phase in the area, as it is observed cross-cutting Matalavaara-type granites in the NW part of the study area. Sample A2163 Mustivaara N is from an approximately 1-m-thick pegmatitic dike that has intruded into quartzite and calc-silicate rocks in the northern part of the PB. The main minerals are quartz, plagioclase and K-feldspar. Tourmaline grains are present as euhedral to subhedral, up to 1-cm-long crystals. The sample also contains euhedral to subhedral garnet grains that may be magmatic in origin. Chlorite, apatite, muscovite and monazite

3.6. A2166 Hosiovaara Sample A2166 Hosiovaara is fine-grained, laminar, glassy, granoblastic sericite quartzite belonging to the Mellajoki Suite. Accessory minerals include tourmaline, zircon and fine-grained opaque minerals.

3.7. A2167 Mäntylaki Sample A2167 Mäntylaki was taken from a fine-grained mica schist of the Martimo Formation (Fig. 2). The rock is multiply folded and contains folded quartz veins and locally cordierite porphyroblasts (Fig. 3a). The main minerals are quartz, biotite and plagioclase and chlorite, muscovite and zircon are found as accessory minerals.

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Fig. 3. Outcrop photographs of (a) folded mica schist of the Martimo Formation (sampling site of A2167 Mäntylaki) and (b) deformed Kierovaara granite pluton (sampling site of A2169 Kierovaara).

3.8. A2168 Louevaara Sample A2168 Louevaara is from a quartzite belonging to the Palokivalo Formation of the Kivalo Group, which represents one of the dominant lithological unit units of the Peräpohja Belt. The rock is gray or reddish and granoblastic and locally displays crossbedding structures. Feldspar and muscovite are found as accessory minerals. 3.9. A2169 Kierovaara Sample A2169 represents the Kierovaara granite pluton having an areal extent of at least 6 × 20 km. It intrudes quartzites of the Mellajoki Suite, just north of the volcanic Väystäjä Formation. The granite is red, K-feldspar-porphyritic and strongly deformed, resembling augen gneiss in the outcrop (Fig. 3b). Potassic feldspar megacrysts, which are 1–4 cm in size, are surrounded by finegrained quartz, plagioclase and biotite. 4. Analytical methods 4.1. U–Pb Zircon grains for U–Pb dating were selected by hand-picking after heavy liquid and magnetic separation of crushed and ground samples. The grains were mounted in epoxy resin and sectioned approximately in half and polished. Back-scattered electron images (BSE) and cathodoluminescence (CL) images were taken of the zircon grains to target the spot analysis sites. U–Pb analyses were performed with a Nu Plasma HR multicollector ICP-MS at the Geological Survey of Finland in Espoo, using a technique very similar to that of Rosa et al. (2009) except that a Photon Machine Analyte G2 laser microprobe was used. Samples were ablated in He gas (gas flows = 0.4 and 0.1 l/min) within a HelEx ablation cell (Müller et al., 2009). Helium aerosol was mixed with Ar (gas flow = 0.8 l/min) prior to entry into the plasma. The gas mixture was optimized daily for achieving the maximum sensitivity. All analyses were carried out in static ablation mode. Ablation conditions were: beam diameter 20 ␮m, pulse frequency 5 Hz, beam energy density 0.55 J/cm2 . A single U–Pb measurement included 30 s of on-mass background measurement, followed by 60 s of ablation with a stationary beam. Masses 204, 206 and 207 were measured in secondary electron multipliers, and 238 in an extra high mass Faraday collector. The geometry of the collector block does not allow simultaneous measurement of 208 Pb and 232 Th. Ion counts were converted and reported as volts by the Nu Plasma time-resolved analysis software. 235 U was calculated from the signal at mass 238

using a natural 238 U/235 U = 137.88. Mass number 204 was used as a monitor for common 204 Pb. In an ICP-MS analysis, 204 Hg mainly originates from the He supply. The observed background countingrate on mass 204 was ca. 1200 (ca. 1.3 × 10−5 V), and has been stable at that level over the last year. The contribution of 204 Hg from the plasma was eliminated by on-mass background measurement prior to each analysis. Age related common lead correction (Stacey and Kramers, 1975) was used when the analysis showed common lead contents above the detection limit. Signal strengths on mass 206 were typically >10−3 V, depending on the uranium content and age of the zircon. Two calibration standards were run in duplicate at the beginning and end of each analytical session, and at regular intervals during sessions. Raw data were corrected for the background, laser induced elemental fractionation, mass discrimination and drift in ion counter gains and reduced to U–Pb isotope ratios by calibration to concordant reference zircons of known age, using protocols adapted from Andersen et al. (2004) and Jackson et al. (2004). Standard zircons GJ-01 (609 ± 1 Ma; Belousova et al., 2006) and an in-house standard A1772 (2712 ± 1 Ma; Huhma et al., 2012) were used for calibration. The calculations were performed offline, using an interactive spreadsheet program written in Microsoft Excel/VBA by T. Andersen (Rosa et al., 2009). To minimize the effects of laser-induced elemental fractionation, the depth-to-diameter ratio of the ablation pit was kept low, and isotopically homogeneous segments of the time-resolved traces were calibrated against the corresponding time interval for each mass in the reference zircon. To compensate for drift in instrument sensitivity and Faraday vs. electron multiplier gain during an analytical session, a correlation of signal vs. time was assumed for the reference zircons. A description of the algorithms used is provided in Rosa et al. (2009). Plotting of the U–Pb isotopic data and age calculations were performed using the Isoplot/Ex 3 program (Ludwig, 2003). All the ages were calculated with 2 errors and without decay constants errors and accordingly, the error ellipses in the figures are at the 2 level. The concordant age offset from ID-TIMS ages does not exceed 0.5%. 4.2. Sm-Nd For whole-rock and garnet Sm–Nd analysis, 120–200 mg of powdered sample was spiked with a 149 Sm-150 Nd tracer. Prior final dissolute garnet was treated following the stepwise dissolution method of DeWolf et al. (1996). This method involves powdering of about 300 mg of separated garnet in a boron carbide mortar for about 30 min and subsequent leaching in hot 6 N HCl for about 6 h, followed by rinsing with deionized water. The sample-spike mixture was dissolved in HF-HNO3 in sealed

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Teflon bombs in an oven at 180 ◦ C for 48 h. Prior to dissolving the residue in 6.2 N HCl, the fluorides were gently evaporated using HNO3 . Conventional cation exchange chromatography was used for separation of the light rare earth elements, and Sm and Nd were separated by a modified Teflon-HDEHP (hydrogen di-ethylhexyl phosphate) method (Richard et al., 1976). Total procedural blank was <0.5 ng for Nd. Isotope ratios were measured on a VG Sector 54 TIMS using Ta-Re triple filaments. Nd isotope ratios were measured in dynamic mode and Sm isotopes in static mode. Nd ratios are normalized to 146 Nd/144 Nd = 0.7219. Based on several duplicate analyses, the error of the 147 Sm/144 Nd is estimated to be better than 0.4%. The long-term average 143 Nd/144 Nd for the La Jolla standard is 0.511850 ± 0.000010 (standard deviation for 220 measurements during the years 1996–2010). Recent analysis on BCR-1 gave Sm = 6.63 ppm, Nd = 28.88 ppm, 143 Nd/144 Nd = 0.512640 ± 0.000010. The εNd was calculated using 147 Sm = 6.54 × 10−12 a−1 , 147 Sm/144 Nd = 0.1966, and 143 Nd/144 Nd = 0.512640 for the present CHUR. TDM was calculated after DePaolo (1981). Plotting and calculations of isotope data were performed using the Isoplot program (Ludwig, 2003). 4.3. Whole-rock geochemistry Major and trace element measurements were performed using X-ray fluorescence (XRF, laboratory method 175X) and inductively coupled plasma mass spectrometry (ICP-MS, laboratory method 308 M) at the Labtium geochemical laboratory. Detailed descriptions of the analytical methods can be found in Rasilainen et al. (2007). 5. Results U–Pb data are presented in Figs. 4 and 5 and listed in Supplementary Appendix 1. Sm–Nd isotopes compositions and whole-rock major and trace-element compositions are shown in Tables 2 and 3, respectively, and Fig. 7 illustrates εNd evolution lines for the studied samples. 5.1. Metasediments A2160 Männikkömaa. Detrital zircon grains from the mica gneiss sample A2160 are partly rounded and subhedral. Totally 64 zircon grains were analyzed (Fig. 4a) of which eight grains showed a high common lead content and were therefore excluded from the age calculations. All zircon grains are probably primarily Archean in age. Some grains display signs of metamorphic overprinting, giving an Archean age from the center of the grain and a Paleoproterozoic age of ca. 1.9 Ga from the rim (Fig. 4b). Garnet-whole rock TIMS analysis gave a Sm–Nd isochron age of 1752 ± 14 Ma (Table 2). The isochron yields an initial εNd (1752 Ma) value of −11.2 and the whole-rock analysis gives a model age (TDM ) of ca. 2860 Ma (Table 3). A2165 Karhurommas. The mica schist sample A2165 yielded abundant zircon. The grains are mostly rounded and show oscillatory zoned cores and thin rims in BSE images. The 207 Pb/206 Pb ages of most zircon grains are Archean, falling mainly in the range of 2.6–2.8 Ga (Fig. 4c). The age distribution is typical for the Archean basement (Fig. 4d). The εNd (1900 Ma) value for whole-rock sample A2165 is −8.4 and the corresponding depleted-mantle model age (TDM ) is ca. 2730 Ma. A2166 Hosiovaara. The quartzite sample A2166 yielded abundant, rounded detrital zircon grains. The 207 Pb/206 Pb ages are Archean, varying between 2.6 Ga and 3.4 Ga (Fig. 4e), with peaks at ca. 2.7 Ga and ca. 2.9 Ga (Fig. 4f). The whole-rock εNd (1900 Ma) value is −10.8 and the corresponding depleted mantle model age (TDM ) is ca. 2990 Ma.

A2167 Mäntylaki. The mica schist sample A2167 produced abundant elongated, partly rounded detrital zircon grains. A total of 49 grains were analyzed, and because four grains contained excess common lead, 45 grains were taken into account in the age calculations (Table 2). U–Pb ages of the zircon grains vary between ca. 1.91 and 3.15 Ga (Fig. 4g). Deviating from the other sedimentary samples, sample A2167 contains a large proportion (66%) of Paleoproterozoic (1.91–2.17 Ga) zircon grains with the rest being Archean (Fig. 4h). Ages between 2.5 Ga and 2.2 Ga seem to be lacking (Fig. 5a). Based on these data, the maximum depositional age for sample A2167 is ca. 1.91 Ga. Sm–Nd analysis of the whole-rock sample gave an εNd (1900 Ma) value of −2.8 with the corresponding depleted mantle model age (TDM ) being ca. 2430 Ma. A2168 Louevaara. Zircon grains from the quartzite sample A2168 are rounded and their grain size varies much. All analyzed 49 zircon grains gave Archean ages, with 207 Pb/206 Pb ages ranging between 2.65 and 3.62 Ga (Fig. 5b). The age distribution (Fig. 4f) is typical for zircon in the Archean crust and consistent with previous zircon data from the Kivalo Group quartzites, which indicate predominance or exclusiveness of Archean detrital zircon grains in these rocks (Perttunen and Vaasjoki, 2001). Whole-rock Sm–Nd analysis produced an εNd (1900 Ma) value of −11.3 and a TDM model age of ca. 3120 Ma.

5.2. Intrusive rocks A2161 Matalavaara. Zircon grains in the alkali-feldspar granite of sample A2161 are elongated and euhedral. The analyzed 28 grains form a reasonably homogeneous population (Fig. 5c), with an upper intercept age of 1775 ± 12 Ma and a lower intercept age of 252 ± 110 Ma (MSWD = 2.6). The former is interpreted as the crystallization age for the granite. Sm–Nd analysis on the wholerock sample gave an initial εNd (1775 Ma) of −6.6 (Table 3). The chondrite-normalized REE pattern is sloping with high (La/Sm)N and moderately high (Gd/Lu)N (Fig. 6). A2162 Matalavaara S. Two different zircon populations were found in sample A2162 that represents alkali-feldspar granitic leucosome in a migmatite. Rounded, anhedral zircon grains gave Archean ages (Fig. 5d), indicating that they are inherited grains from the host metasediment containing Archean detrital zircon. On the other hand, subhedral to euhedral zircon grains yielded an average 207 Pb/206 Pb age of 1793 ± 12 Ma, which is considered as the best estimate for the crystallization age of the leucosome (Fig. 5d). Three zircon analyses were excluded from the age calculation due to their high common lead content. The whole-rock Sm–Nd analysis gave an initial εNd (1793 Ma) value of −6.6 (Table 3). The chondrite-normalized REE pattern displays slightly fractionated LREE, whereas HREE are essentially flat. A2163 Mustivaara N and A2164 Mustivaara S. Of these two tourmaline-bearing pegmatitic granite samples, only sample A2163 yielded a small amount of zircon. The grains are elongated and partly rounded. With the exception of one grain, they gave Archean ages between 2.6 and 3.1 Ga (Fig. 5e). An Archean age for sample A2163 can be ruled out, as it was collected from a dike that cross-cuts Paleoproterozoic metasediments. Moreover, similar tourmaline granite dikes are observed to cross-cut Matalavaara-type granites elsewhere in the study area. The old zircon grains most probably represent exotic grains that were captured from a quartzite country rock during the intrusion of the granitic dike. One high-U zircon grain yielded a 207 Pb/206 Pb age of ca. 1761 Ma, which may represent the crystallization age of the pegmatite. Both tourmaline granites show flat chondritenormalized REE patterns (Fig. 6), consistent with the relatively high measured 147 Sm/143 Nd ratios. Assuming a crystallization age of

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Fig. 4. U–Pb Concordia diagrams, histograms and probability distribution of zircon isotopic data.

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Fig. 5. U–Pb Concordia diagrams and probability distribution of zircon isotopic data.

1760 Ma, initial εNd values of −11.4 and −11.3 can be calculated (Table 3). A2169 Kierovaara. Zircon grains of sample A2169, representing a granitic pluton that intrudes quartzites of the Mellajoki Suite, are euhedral and magmatic in appearance. The amount of common lead is very low. Zircon analyses gave a concordant U–Pb age of 1989 ± 6 Ma, which is considered the crystallization age of the granite (Fig. 5f). Whole-rock Sm–Nd analysis gave an initial εNd (1989 Ma) value of −3.9. The initial εNd value is more radiogenic compared to the initial values obtained for younger granites analyzed in this work, but the calculated Nd isotopic evolution of

the Kierovaara granite fits rather well with that of the 1.78 Ga Matalavaara granite (A2161) (Fig. 7). The granite shows a fractionated chondrite-normalized REE pattern with a moderate Eu anomaly (Fig. 6). 6. Discussion 6.1. Metasediments Two quartzite samples representing the Mellajoki Suite (A2166) and the Palokivalo Formation (A2168) were selected for dating

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Table 2 TIMS Sm–Nd isotopic data. Sample

Location

Rock type

Sm (ppm)

Nd

147

Sm/144 Nd

A2160 A2160 grt A2161 A2162 A2163 A2164 A2165 A2166 A2167 A2168 A2169

Männikkömaa Männikkömaa Matalavaara Matalavaara S Mustivaara N Mustivaara S Karhurommas Hosiovaara Mäntylaki Louevaara Kierovaara

Mica gneiss Garnet Granite Granite Granite Granite Mica schist Quartzite Mica schist Quartzite Granite

5.52 0.37 2.80 3.57 2.01 0.84 10.49 0.54 5.57 1.07 4.78

31.83 0.32 20.69 14.41 5.40 1.93 63.94 3.03 28.81 5.59 30.46

0.1048 0.6926 0.0819 0.1496 0.2254 0.2638 0.0992 0.1081 0.1168 0.1156 0.0948

143

Nd/144 Nd

0.511005 0.517778 0.510966 0.511650 0.512397 0.512847 0.510996 0.510982 0.511499 0.511052 0.511106

2se

T (Ma)

ε (T)

T-DMa (Ma)

0.000010 0.000046 0.000010 0.000018 0.000011 0.000031 0.000010 0.000014 0.000011 0.000016 0.000010

1900 1700 1775 1793 1780 1780 1900 1900 1900 1900 1989

−9.6 −8.0 −6.6 −8.5 −11.4 −11.4 −8.4 −10.8 −2.8 −11.3 −4.0

2864 2400 3386

2730 2990 2429 3116 2483

Error in 147 Sm/144 Nd is 0.4%. Sample coordinates shown in Appendix 1. a Depleted mantle model age after DePaolo (1981).

Table 3 Major and trace element analyses of metasedimentary and granitic rocks. A2160

A2161

A2162

A2163

A2164

A2165

A2166

A2167

A2168

A2169

SiO2 (wt%) TiO2 Al2 O3 Fe2 O3tot MnO MgO CaO Na2 O K2 O P2 O5 S

60.2 0.84 18.3 8.32 0.08 2.39 2.45 4.53 2.43 0.09 0.01

74.5 0.09 14.0 1.60 b.d. 0.22 0.67 3.56 5.14 0.03 b.d.

73.7 0.06 14.7 0.88 b.d. 0.29 0.55 3.16 6.47 0.07 b.d.

76.2 0.02 13.5 0.76 0.09 0.09 0.46 3.09 5.55 0.15 b.d.

76.0 0.02 14.9 1.15 0.01 0.23 0.40 6.25 0.87 0.09 b.d.

53.2 0.88 20.1 16.6 0.17 3.06 1.28 1.16 2.39 0.35 0.03

n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.

64.5 0.90 15.8 8.70 0.08 4.11 0.63 1.15 2.15 0.16 0.23

88.9 0.07 5.20 1.37 0.02 0.88 0.19 0.07 2.16 0.06 b.d.

71.7 0.34 14.1 3.23 0.04 0.83 1.40 3.14 4.85 0.09 0.04

Total Cr (ppm) Ni Cu Zn Ga Rb Sr Y Zr Ba Pb Co Zr Hf Nb Ta Rb Sc V U Th Y La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

99.6 150 70 b.d. 80 30 130 330 19 210 710 b.d. 24.2 156 4.23 10.9 1.21 105 18.1 131 3.36 14.8 26.5 36.6 74.1 8.5 32.1 5.82 1.47 5.58 0.83 5.06 1.04 2.96 0.44 2.71 0.4

99.8 b.d. b.d. b.d. b.d. 20 180 90 15 120 460 30 1.42 147 4.41 3.66 0.36 174 3.05 13.2 4.94 65.1 9.61 43.3 86.3 7.31 22.9 3.15 0.63 2.61 0.33 1.74 0.33 0.96 0.14 0.83 0.12

99.9 20 b.d. b.d. b.d. b.d. 160 100 46 30 600 50 0.95 34.7 1.37 3.05 0.39 143 3.18 8.34 19.4 25.8 43.7 18.8 37.4 4.23 15.4 3.79 0.48 4.48 0.93 6.95 1.54 4.98 0.78 4.93 0.71

99.9 b.d. b.d. b.d. b.d. b.d. 590 b.d. 36 20 30 b.d. b.d. 30.2 1.47 32.9 8.56 519 5.47 7.67 3.47 7.84 13.1 4.88 10.5 1.38 5.82 2.15 0.16 2.09 0.41 2.47 0.4 1.26 0.25 2.01 0.31

99.9 b.d. b.d. b.d. b.d. b.d. 70 b.d. 11 20 b.d. b.d. 0.97 30.6 1.12 6.57 0.99 65.4 5.28 6.39 2.47 4.52 7.16 2.36 5.59 0.65 2.33 0.98 b.d. 0.9 0.21 1.4 0.23 0.81 0.19 1.56 0.25

99.2 170 90 20 100 30 120 70 32 110 420 b.d. 30.7 121 3.36 12.2 1.27 107 31.4 196 5.56 21.5 30.8 70.8 146 16.6 63 10.7 1.77 9.17 1.26 6.79 1.25 3.54 0.49 3.09 0.46

n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. 0.64 42.3 1.13 0.33 b.d. 15.4 1.89 16.7 0.26 1.71 6.18 4.32 8.91 0.92 3.35 0.64 0.19 0.79 0.16 1.14 0.23 0.71 0.1 0.52 b.d.

98.4 250 130 160 100 20 90 70 27 180 310 b.d. 24.2 172 4.75 10.6 1.09 80.9 21.6 129 3.22 11.3 24.5 33.4 67.3 7.83 30 5.88 1.01 5.74 0.88 5.14 0.98 2.74 0.4 2.33 0.35

98.9 40 b.d. b.d. b.d. b.d. 50 b.d. 12 100 210 b.d. 0.81 100 2.5 1.48 b.d. 47.7 2.7 14.5 0.55 3.54 4.96 6.38 12.9 1.45 5.7 1.07 0.25 1.16 0.16 0.94 0.19 0.56 b.d. 0.52 b.d.

99.7 30 b.d. b.d. 30 20 190 120 26 170 780 30 6.3 169 4.68 12.2 1.42 159 7.5 37.8 6.27 39.6 17 55.9 97.5 9.15 30 4.67 0.6 4.25 0.62 3.57 0.65 1.84 0.26 1.69 0.24

Sample coordinates shown in Appendix 1. Major elements and trace elementa from Cr to Pb determinated by XRF, trace elements from Co to Lu determined by ICP-MS. b.d., below detection limit; n.d., not determined.

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Fig. 6. Chondrite-normalized REE patterns for granitoids from the northern part of the Peräpohja Belt. Normalizing values after Sun and McDonough (1989). A2161 alkali-feldspar granite, A2162 granitic leucosome in migmatite, A2163 and A2164 tourmaline granites, A2169 K-feldspar-porphyritic, deformed granite. The gray field represents 1.98 Ga felsic tuffs of the Korkiavaara Formation having an A-type granitic chemical composition (Hanski et al., 2005).

4

ƐNd

0

DM (DePaolo)

CHUR

-4

A2169 Kierovaara

A2161 Matalavaara

-8

A2162 Matalavaara S

A2163 Mustivaara N

-12

A2164Mustivaara S

-16 1500

1700

1900

2100

2300

2500

2700

2900

Age (Ma) Fig. 7. εNd vs. age diagram showing initial εNd values for granitoid samples. Evolution lines are shown for metasedimentary samples. CHUR, Chondritic uniform reservoir. DM, Depleted mantle evolution after DePaolo (1981).

detrital zircon populations in these metasediments. Perttunen and Hanski (2003) previously regarded the Mellajoki Suite as a part of the Central Lapland Granitoid Complex based on the strong deformation seen in the rocks compared with the quartzites of the Kivalo Group. Perttunen and Vaasjoki (2001) reported Archean multi-grain TIMS ages for detrital zircons from the Palokivalo Formation but no previous age data are available from the quartzite unit in the Mellajoki Suite. In this study, the quartzite samples from both rock units revealed rounded detrital zircon populations that are exclusively Archean in age, though the age distribution of the detrital zircon grains vary between two samples. In the Mellajoki Suite sample (A2166), two largest zircon populations occur at ca. 2.7 Ga and 2.9 Ga with a smaller peak at ca. 3.1 Ga (Fig. 4f). In the sample from the Palokivalo Formation (A2168), the largest zircon population also occurs at ca. 2.7 Ga but the 2.9 Ga peak is missing. A minimum age of ca. 2.22 Ga for the Palokivalo Formation has been determined from the differentiated mafic sills belonging to the gabbro-wehrlite (GWA) association (Perttunen and Vaasjoki, 2001; Hanski et al., 2010). Mafic dikes have also been reported to

cut the Mellajoki Suite by Perttunen and Hanski (2003), but no age data are available from these dikes. Recent drilling in the Rompas area has revealed evidence for the presence of differentiated sills in the area, with characteristics typical of the 2.22 Ga intrusions (Huttu, 2014). Based on these observations and our zircon data, it is plausible to suggest that the Mellajoki Suite and Palokivalo Formation represent the same stage of sedimentation in the PB. The varying degree of deformation and the slightly different composition of the rock units, i.e. the more arkosic mineralogy of the sample A2168 Louevaara compared with the orthoquartzitic sample A2166 Hosiovaara probably only indicate that the samples were collected from different stratigraphic levels. As was mentioned before, the Kalevian rock successions are divided into the autochthonous-parautochthonous Lower Kaleva and dominantly allochthonous Upper Kaleva (Kontinen, 1987). Until recently, the maximum age for the deposition of the Lower Kalevian interbedded metagraywackes and metapelitic rocks has been thought to be ca. 2.1 Ga, based on the age of the underlying Jatulian mafic and felsic volcanic rocks (e.g., Huhma, 1986; Pekkarinen and Lukkarinen, 1991; Lahtinen et al., 2010), or 1.97 Ga, the age of the youngest mafic dike generation, which does not reach the stratigraphic level the Kalevian rocks (Kohonen, 1995). On the other hand, based on the close association of Upper Kalevian metasediments with rocks of the ca. 1.95 Ga Jormua Ophiolite in the Kainuu Schist Belt and the Outokumpu ophiolitic serpentinites in the North Karelia Belt, at least part of the Upper Kaleva is considered to have deposited on oceanic crust and therefore must be considerably younger that the Lower Kalevian rocks. Consistent with this interpretation, in situ dating of detrital zircon grains in Upper Kalevian metasediments has revealed the presence of a significant Paleoproterozoic component indicating maximum depositional ages of 1.95–1.92 Ga (Claesson et al., 1993; Lahtinen et al., 2010). In the absence of obvious ophiolitic rocks in the Peräpohja Belt, the mica schists of the Martimo Formation have been regarded as Lower Kalevian (Laajoki, 2005). This is also compatible with the limited earlier zircon data (Perttunen and Vaasjoki, 2001; Hanski et al., 2005). However, our new results show that this view needs to be reconsidered. In our sample A2167 Mäntylaki, 66% of the detrital zircon grains are Paleoproterozoic in age and 97% of these zircon grains fall between ca. 1.9 Ga and 2.1 Ga. The lack of detrital zircon populations between 2.5 and 2.2 Ga in sample A2167 is similar to what Lahtinen et al. (2010) reported for Upper Kalevian samples from the North Karelia Belt. They also reported whole-rock εNd values (at 1.9 Ga) from −0.6 to −2.8 for these metasediments, which fit well with the whole-rock εNd value (at 1.9 Ga) of −2.8 obtained for our sample. Recently, Lahtinen et al. (2015) conducted a LAICP-MS study of the detrital zircon separate from the same mica schist sample (A720) from the Martimo Formation that was earlier studied by Perttunen and Vaasjoki (2001) using bulk zircon TIMS analysis. Perttunen and Vaasjoki (2001) analyzed one bulk zircon fraction, which yielded an age of ca. 2575 Ma. Instead, about half of the zircon grains analyzed by Lahtinen et al. (2015) are Paleoproterozoic in age, indicating a maximum depositional age of 1.95 Ma. A720 is thus the second sample (in addition to our A2167) from the Peräpohja Belt, recording a time of sedimentation that fits better with that of the Upper Kaleva than the Lower Kaleva. Among the observations that have led to the conclusion that the graywackes and black shales of the Martimo Formation belong to the Lower Kaleva is the occurrence of pillow basalts and felsic porphyries of the Väystäjä Formation in a close spatial association with Martimo Formation metasediments (see Fig. 2). The U–Pb age of 2050 ±8 Ma of the felsic porphyries in the Väystäjä Formation (Perttunen and Vaasjoki, 2001) demonstrates that the environment where the volcanic or subvolcanic porphyries were emplaced is at least ca. 2050 Ma in age and hence stratigraphically belongs to the Lower Kaleva rather than to the Upper Kaleva.

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On the other hand, dolomites associated with the pillow lavas of the Väystäjä Formation record a normal oceanic ı13 C composition attesting that their deposition took place after the c. 2.06 Ga, which is regarded as the age for the end of the Lomagundi-Jatuli carbon isotope excursion (Melezhik et al., 2007; Martin et al., 2013). Our detrital zircon age data clearly indicate that at least some of the Martimo Formation sediments were deposited much later than the Väystäjä Formation. One solution to this problem is to regard the Väystäjä Formation as an allochthonous unit or even a piece of overthrust oceanic crust, which would be consistent with the chemical and isotopic composition of its metalavas and carbonate metasediments. In summary, in situ dating shows that some of the mica schist and mica gneiss samples from the Peräpohja Belt contain solely Archean zircon populations (A2160 and A2165; also A1571 of Hanski et al., 2005), whereas some other samples (A2167 and A720) contain mixed Archean and Paleoproterozoic populations, with the youngest grains being 1.95–1.91 Ga in age. It thus seems that at least two distinct Kalevian metasedimentary units exist in the Peräpohja Belt, implying that there potentially exists an unconformity within the Kalevian sequence. The Taivalkoski conglomerate (Perttunen, 1985) in the SW part of the belt is one candidate for the manifestation of such a great break in sedimentation. The other alternative is that all metapelites of the Martimo Formation are Upper Kalevian in age and the unconformity or tectonic contact is located between the Kivalo and Paakkola Groups. When comparing detrital zircon age distributions in the Martimo Formation and other Kalevian formations in Finland, similarities appear. In their study of metasediments from the North Karelia Belt, Lahtinen et al. (2010) determined in situ ages of a detrital zircon separate from one Lower Kalevian metagraywacke, yielding solely Archean ages. This result together with earlier bulk zircon TIMS data and geochemical and Nd isotopic characteristics of Lower Kalevian metasediments was used to formulate the conclusion that the Lower Kalevian sediments had a mixed source composed of Archean basement rocks and Paleoproterozoic (Karelian) cover material including 2.1 Ga mafic lavas and dikes. In contrast, the provenance of Upper Kalevian zircon grains dated at 1.95–1.92 Ga was interpreted by Lahtinen et al. (2010) to have been the Himalaya-type Lapland-Kola Orogen in northern Fennoscandia. This view with the Lower and Upper Kalevian rocks having different depositional ages and sources of detritus has, however, recently been questioned when additional isotopic data have become available from Lower Kalevian rocks in the North Karelia Belt. Lahtinen et al. (2013) published new detrital zircon data from several lithological assemblages in the Höytiäinen basin, which have earlier been assigned to the Lower Kaleva (Lahtinen et al., 2010), as they differ from typical ophiolite-bearing allochthonous sedimentary strata in the region (Kohonen, 1995). The zircon ages indicated relatively young maximum depositional ages of 1.92–1.91 Ga for these metasediments. Lahtinen et al. (2013) interpreted the studied metagraywackes as foredeep to foreland basin sediments with their deposition related to the continent-arc/continent collision between the Archean craton and a Paleoproterozoic microcontinent-arc collage. From the discussion above, we can conclude that the Kalevian rocks seem to represent deep-water metasediments, for which diverse ages of deposition and geotectonic environments have been proposed, with the rocks being regarded as <1.97 Ga continental margin or intracratonic deposits, <1.95 Ga pelagic ocean floor sediments, and <1.91–1.94 Ga foredeep/foreland basin sediments. In addition, the Central Lapland Greenstone belt contains phyllites and black shales that are older than or equal to ca. 2.06 Ga in age, which is the age of the Kevista mafic-ultramafic intrusion emplaced into phyllitic sediments (Mutanen and Huhma, 2001). Hence, the division of the Kalevian rocks into two main units, the Lower and

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Upper Kaleva, has become problematic. More research and particularly in situ dating of detrital minerals are needed to clarify the age distribution, provenance and geotectonic settings of the Kalevian rocks in Finland. 6.2. Intrusive rocks and migmatites With respect to the Svecofennian orogenic events, the Paleoproterozoic granitic magmatism in Finland has been divided into several stages: (1) preorogenic (1.95–1.91 Ga), (2) synorogenic divided into synkinematic (1.89–1.87 Ga) and postkinematic (1.88–1.86 Ga) phases, (3) lateorogenic (1.84–1.80 Ga), and (4) postorogenic (1.81–1.77 Ga) (Nironen, 2005). However, the division is not wholly applicable to the area NE of the Raahe–Ladoga zone that separates the Svecofennian area sensu stricto from the Archean Karelian terrain and its Paleoproterozoic cover rocks. In the Peräpohja Belt and its surrounding areas, the synorogenic stage is represented by the Haaparanta Suite monzonitic plutons (1.89–1.86 Ga; Perttunen and Vaasjoki, 2001) and the postorogenic stage by small granitoid stocks such as the Rovaniemi granite dated at ca. 1.77 Ga (Lauerma, 1982). In this study, three (A2161, A2162, A2163) of the four dated granite or granitic leucosome samples gave ages between 1.77 and 1.79 Ga, thus belonging to the period of widespread granitoid magmatism and migmatization that took place in northern Finland at that time, but which cannot be classified as postorogenic, as most granitoids of this age were emplaced during tectonic activity and commonly show a foliation (Ahtonen et al., 2007; Lauri et al., 2012). Strongly negative initial εNd values, ranging from −6.6 to −11.4, indicate that these granites had a major Archean component in their source. Together with inherited Archean zircon grains in sample A2162, these findings support the presence of Archean basement at the time of granite emplacement and are in line with previous Sm–Nd isotopic studies of granites emplaced in the area of the Karelian craton (e.g., Huhma, 1986). The ca. 1.8 Ga rims around the Archean zircon grains that are found in the migmatite sample A2160 further support the concept of a migmatization event at 1.79–1.77 Ga, as proposed by Ahtonen et al. (2007). On the other hand, the Sm–Nd age of 1752 ± 14 Ma that was determined for a whole rock-garnet pair from sample A2160 reflects cooling of the crust down to the closure temperature of garnet, which has been estimated to be between 600 ◦ C (Metzger et al., 1992) and 700 ◦ C (Hensen and Zhou, 1995). Further studies are needed to confirm if any of the dates mentioned above are related to the formation of the gold mineralization in the study area. One intrusion, the Kierovaara porphyritic granite pluton, which intrudes quartzites of the Mellajoki Suite, yielded a clearly older, concordant zircon age of 1989 ± 6 Ma. The Kierovaara granite is obviously pre-orogenic in terms of its timing and furthermore, the age is outside the range of ages listed above for the main Paleoproterozoic stages of granitic magmatism. The Kierovaara granite thus seems to represent a previously unrecognized phase of felsic plutonism in northern Finland and even in the whole Fennoscandian Shield (cf. Nironen, 2005; Hanski et al., 2001; Hanski and Melezhik, 2012). Nevertheless, an indication of the former presence of such a magmatic phase is recorded by detrital zircon grains with an age of ca. 1.98 Ga that have been observed in Upper Kalevian and Svecofennian metasediments (e.g., Huhma et al., 1991; Claesson et al., 1993; Hanski et al., 2005; Lahtinen et al., 2010, 2015). In addition, one piece of evidence for the existence of felsic magmatism that is approximately contemporaneous with the Kierovaara granite comes from a zircon population dated by Hanski et al. (2005) from the so-called “arkosites” of the Korkiavaara Formation that occur in the northern part of the Peräpohja Belt. These felsic rocks have typical chemical characteristics of A-type granites and are interbedded with amphibolites in scales ranging from a few millimeters to tens of meters. Hanski et al. (2005) discussed two alternative

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hypotheses for the generation of the felsic rocks of the Korkiavaara Formation. A single zircon population found in three different locations together with geochemical data indicates that these rocks are either homogeneous epiclastic sediments derived from a single source of A-type granites or they represent tuffs of coeval mafic and felsic magmas. In any case, the observations by Hanski et al. (2005) demonstrate an event of A-type ca. 1.98 Ga granitic magmatism that can potentially be linked to the mantle plume-related maficultramafic magmatism of the Pechenga and Onega regions in the eastern part of the shield (Walker et al., 1997; Puchtel et al., 1998). However, even though being of the same age as the Korkiavaara Formation, the Kierovaara granite is lower in REE (Fig. 6) and HFSE, with no clear A-type chemical signature. The possible relationship between the Kierovaara granite and the coeval Korkiavaara Formation needs more study. A zone of quartz-feldspar gneiss have been reported just south of the Kierovaara granite (Fig. 2), having indistinguishable chemistry compared that of the Korkiavaara Formation felsic rocks (Perttunen and Hanski, 2003), but their contact is most probably tectonic. Recent dating by Lahtinen et al. (2015) shows that this felsic unit has an age of ca. 1.98 Ga, confirming its connection with the Korkiavaara Formation rocks. Taking into account the new zircon data from the Martimo Formation, we can conclude that the magma generation and emplacement of the Kierovaara granite took place during a late stage of deposition of the Karelian sedimentary-volcanic sequence with no apparent link to orogenic events. The strong deformation of the granite is compatible with its zircon age. Though being geochronologically anomalous, the Kierovaara granite is not the only representative of abnormally old granites that are found around the Central Lapland Granite Complex. Rastas et al. (2001) reported on the occurrence of ca. 2.13 Ga granites at the northern margin of the complex, and later using in situ zircon dating, Ahtonen et al. (2007) and Lauri et al. (2012) confirmed the presence of roughly coeval granodiorites and granites in the eastern part of the complex. The mechanism causing extensive crustal melting in a non-orogenic, extensional geotectonic setting at ca. 2.11–2.13 Ga is not clear, but could be related to mafic magma underplating in lower crust. However, it is worth of mentioning that at least the 2.1 Ga extrusive and subvolcanic mafic magmas in the Peräpohja Belt, as represented by the Jouttiaapa Formation, avoided significant interaction with continental crust (Huhma et al., 1990; Kyläkoski et al., 2012). The Kierovaara granite seems to be 100 Ma younger than the above mentioned “old” granitoids and therefore unrelated to the geological events that produced the latter.

7. Conclusions Based on our single zircon U–Pb and whole-rock Sm–Nd data, we can list the following conclusions:

1) Granites and migmatites in the study area represent at least two distinct magmatic phases that were emplaced at 1.99 Ga and 1.79–1.77 Ga, with the younger event representing the main stage of the generation of the Central Lapland Granitoid Complex, whereas the older event is a previously unrecognized episode of felsic plutonism in northern Finland. 2) Archean basement rocks existed under the Paleoproterozoic supracrustal formations during the granitic magmatism both at 1.77–1.79 Ga and 1.99 Ga. 3) Two quartzitic units, the Jatulian Palokivalo Formation of the Kivalo Group and the lithodemic Mellajoki Suite, occurring in different parts of the Peräpohja Belt, display Archean detrital zircon populations and most probably represent sedimentary units that can be correlated with each other.

4) The dates of the youngest detrital zircon grains demonstrate that the maximum depositional age of some parts of the Martimo Formation is ca. 1.91, whereas the synorogenic magmatism of the Haaparanta suite places the upper age limit at ca. 1.88 Ga. This time interval is much younger than previously thought for Kalevian rocks in the Peräpohja belt. Together with the recently published isotopic data from the North Karelia Belt, our data stress the importance of performing more in situ dating of detrital minerals from Kalevian rocks in Finland to clarify their age distribution, provenance and geotectonic setting. Acknowledgments Tuula Hokkanen and Arto Pulkkinen are thanked for assisting in the isotope laboratory of the Geological Survey of Finland. This work was supported by Mawson Resources Ltd. which is gratefully acknowledged. We thank Fernando Corfu and one anonymous reviewer who gave useful comments and suggestions that significantly improved this paper. Appendix A. Supplementary data Supplementary data associated with this article can be found, in the online version, at http://dx.doi.org/10.1016/j.precamres.2015. 05.018 References Andersen, T., Griffin, W.L., Jackson, S.E., Knudsen, T.-L., Pearson, N.J., 2004. MidProterozoic magmatic arc evolution at the southwest margin of the Baltic Shield. Lithos 73, 289–318. Ahtonen, N., Hölttä, P., Huhma, H., 2007. Intracratonic Palaeoproterozoic granitoids in northern Finland: prolonged and episodic crustal melting events revealed by Nd isotopes and U–Pb ages on zircon. Bull. Geol. Soc. Finland 79, 143–174. Belousova, E.A., Griffin, W.L., O’Reilly, S.Y., 2006. Zircon crystal morphology, trace element signatures and Hf isotope composition as a tool for petrogenetic modeling: examples from Eastern Australian granitoids. J. Petrol. 47, 329–353. Claesson, S., Huhma, H., Kinny, P.D., Williams, I.S., 1993. Svecofennian detrital zircon ages – implications for the Precambrian evolution of the Baltic Shield. Precambrian Res. 64, 109–130. DePaolo, D.J., 1981. Neodymium isotopes in the Colorado Front Range and crustmantle evolution in the Proterozoic. Nature 291, 684–687. DeWolf, C.P., Zeissler, C.J., Halliday, A.N., Mezger, K., Essene, E.J., 1996. The role of inclusions in U–Pb and Sm–Nd garnet geochronology: stepwise dissolution experiments and trace uranium mapping by fission track analysis. Geochim. Cosmochim. Acta 60, 121–134. Hanski, E., 2001. History of stratigraphical research in northern Finland. Geol. Surv. Finland 33, 15–43 (special paper). Hanski, E., 2012. Evolution of the Palaeoproterozoic (2. 50–1.95 Ga) non-orogenic magmatism in the eastern part of the Fennoscandian Shield. In: Melezhik, V.A., Prave, A.R., Hanski, E.J., Fallick, A.E., Lepland, A., Kump, L.R., Strauss, H. (Eds.), Reading the Archive of Earth’s Oxygenation. The Palaeoproterozoic of Fennoscandia as Context for the Fennoscandian Arctic Russia – Drilling Early Earth Project, vol. 1. Springer-Verlag, Berlin/Heidelberg, pp. 179–245. Hanski, E.J., Melezhik, V.A., 2012. Litho- and chronostratigraphy of the Karelian formations. In: Melezhik, V.A., Prave, A.R., Hanski, E.J., Fallick, A.E., Lepland, A., Kump, L.R., Strauss, H. (Eds.), Reading the Archive of Earth’s Oxygenation. Volume 1: The Palaeoproterozoic of Fennoscandia as Context for the Fennoscandian Arctic Russia – Drilling Early Earth Project. Springer-Verlag, Berlin/Heidelberg, pp. 39–110. Hanski, E., Huhma, H., Vaasjoki, M., 2001. Geochronology of northern Finland: a summary and discussion. Geol. Surv. Finland Spec. Pap. 33, 255–279. Hanski, E., Huhma, H., Perttunen, V., 2005. SIMS, Sm–Nd isotopic and geochemical study of an arkosite-amphibolite suite, Peräpohja Schist Belt: evidence for 1.98 Ga A-type felsic magmatism in northern Finland. Bull. Geol. Soc. Finland 77, 5–29. Hanski, E., Huhma, H., Vuollo, J., 2010. SIMS zircon ages and Nd isotope systematics of the 2.2 Ga mafic intrusions in northern and eastern Finland. Bull. Geol. Soc. Finland 82, 31–62. Hensen, B.J., Zhou, B., 1995. Retention of isotopic memory in garnets partially broken down during an overprinting granulite-facies metamorphism: implications for the Sm–Nd closure temperature. Geology 23, 225–228. Hölttä, P., Väisänen, M., Väänänen, J., Manninen, T., 2007. Paleoproterozoic metamorphism and deformation in Central Finnish Lapland. Geol. Surv. Finland 44, 109–120 (special paper). Hölttä, P., Huhma, H., Lahtinen, R., Nironen, M., Perttunen, V., Vaasjoki, M., Väänänen, J., 2003. Introduction: modelling of orogeny in northern Fennoscandia. In:

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