Paleoproterozoic magmatism across the Archean

Lithos 262 (2016) 507–525

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Paleoproterozoic magmatism across the Archean-Proterozoic boundary in central Fennoscandia: Geochronology, geochemistry and isotopic data (Sm–Nd, Lu–Hf, O) Raimo Lahtinen a,⁎, Hannu Huhma a, Yann Lahaye a, Stefanie Lode b, Suvi Heinonen a, Mohammad Sayab a, Martin J. Whitehouse c a b c

Geological Survey of Finland, P.O. Box 96, FI-02151 Espoo, Finland Memorial University of Newfoundland, 300 Prince Philip Drive, A1C 2H1 St. John's, NL, Canada Department of Geosciences, Swedish Museum of Natural History, Box 50007, SE-104 05 Stockholm, Sweden

a r t i c l e

i n f o

Article history: Received 7 April 2016 Accepted 12 July 2016 Available online 25 July 2016 Keywords: Paleoproterozoic Fennoscandia Oxygen isotopes U–Pb Sm–Nd Lu–Hf

a b s t r a c t The central Fennoscandia is characterized by the Archean-Proterozoic (AP) boundary and the Central Finland Granitoid Complex (CFGC), a roundish area of approximately 40,000 km2 surrounded by supracrustal belts. Deep seismic reflection profile FIRE 3A runs across these units, and we have re-interpreted the profile and crustal evolution along the profile using 1.92–1.85 Ga plutonic rocks as lithospheric probes. The surface part of the profile has been divided into five subareas: Archean continent (AC) in the east, AP, CFGC, boundary zone (BZ) and the Bothnian Belt (BB) in the west. There are 12 key samples from which zircons were studied for inclusions and analyzed (core-rim) by ion probe for U–Pb dating and oxygen isotopes, followed by analyzes for Lu–Hf by LA–MC–ICP–MS. The AC plutonic rocks (1.87–1.85 Ga) form a bimodal suite, where the proposed mantle source for the mafic rocks is 2.1–2.0 Ga metasomatized lower part of the Archean subcontinental lithospheric mantle (SCLM) and the source for the felsic melts is related plume-derived underplated mafic material in the lower crust. Variable degrees of contamination of the Archean lower crust have produced “subduction-like” Nb–Ta anomalies in spidergrams and negative εNd (T) values in the mafic-intermediate rocks. The felsic AC granitoids originate from a low degree melting of eclogitic or garnet-bearing amphibolites with titanite ± rutile partly prevailing in the residue (Nb–Ta fractionation) followed by variable degree of assimilation/melting of the Archean lower crust. The AP plutonic rocks (ca. 1.88 Ga) can be divided into I-type and A-type granitoids (AP/A), where the latter follow the sediment assimilation trend in ASI diagram, have high δ18O values (up to 8‰) in zircons and exhibit negative Ba anomalies (Rb–Ba–Th in spidergram), as found in sedimentary rocks. A mixing/assimilation of enriched mantle-derived melts with melts from already migmatized sedimentary rocks ± amphibolites is proposed. The CFGC is characterized by both I-type and A-type (CFGC/A) intermediate and felsic granitoids. The I-type granitoids are divided into two groups at ≥1885 Ma and ≤1882 Ma, where the latter overlap in age with the CFGC/A granitoids. Both I-type CFGC and CFGC/A granitoids are interpreted to have formed from mixing of Paleoproterozoic SCLM-derived melts with crustal melts from hydrous and dry intermediate-felsic igneous sources, respectively. The geochemistry, dominantly δ18O values below 6.5‰ in zircons and TDM (2.11– 2.42 Ga) of the CFGC granitoids favor the occurrence of older crust (ca. 2.1–2.0 Ga) in their genesis. The BZ granitoids are similar in age but more juvenile with TDM ages between 1.94 Ga and 2.16 Ga. The 1.92 Ga granodiorite in the BB is correlated with juvenile gneissic tonalites and granodiorites found from the AP boundary. We suggest that the present high-velocity lower crust under the CFGC is composed of melt-extracted granulites (crustal source age ≥ 2.0 Ga) and mafic cumulates which both formed during 1.90–1.88 Ga arc magmatism. The ≤ 1.88 Ga stage represents the end of compression/transpression and is followed by 1.87–1.86 Ga buckling, forming the Bothnian Oroclines. © 2016 Elsevier B.V. All rights reserved.

1. Introduction

⁎ Corresponding author. E-mail address: raimo.lahtinen@gtk.fi (R. Lahtinen).

http://dx.doi.org/10.1016/j.lithos.2016.07.014 0024-4937/© 2016 Elsevier B.V. All rights reserved.

Granitoids are most abundant rocks in the upper continental crust (Wedepohl, 1991) and characteristic for the Precambrian shield areas (e.g., Koistinen et al., 2001). They form in diverse tectonic settings and

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R. Lahtinen et al. / Lithos 262 (2016) 507–525

their geochemical composition has been used as an indicator of distinct tectonic processes (e.g., Pearce et al., 1984). Granitoid magmas can form as differentiation products of mantle-derived mafic magmas (e.g., Fowler et al., 2008) or via crustal melting (Brown, 2013). Differentiation of mafic magmas to felsic composition needs a large component of mafic cumulates (e.g., Annen et al., 2015) and purely crustal-derived melts should be felsic and peraluminous (Patiño Douce, 1999). It appears that the most common intermediate metaluminous granitoids are a mixture of crustal (igneous-sedimentary) and mantle (depleted-enriched) source components. Especially, under thick continental crust, e.g. in continental margin settings, the MASH (Melting, Assimilation, Storage, Homogenization) zones develop in lower crust and lead to vast volumes of dominantly intermediate magmas as batholiths in upper crust (Hildreth and Moorbath, 1988). Brown (2013) accepted the latter mechanism for arc magmas but considers it less applicable to the generation of granites associated with collisional orogenesis. Thicker, older crust also promotes larger amounts of contamination/assimilation of the ascending mafic magmas and thus, the mafic magmas and their differentiation products intruding to middle- and upper-crust may not represent the mantle composition but, instead, show crustal trace element fingerprints. The availability of both heat and water limit the extent of crustal melting (Aranovich et al., 2014; Weinberg and Hasalová, 2015) and indicates that either hydrous fluids or/and large amount of mafic magmas are needed to create large volumes of crustal melts. Subduction-related mafic magmas are hydrous (e.g., Grove et al., 2002) and can exsolve H2O from residual melt possibly triggering water saturated melting of crust. Within-plate mafic magmas are typically anhydrous and dehydration melting or external sources of hydrous fluid are needed to produce crustal melts. Restite unmixing and incomplete magma mixing are seen as enclaves but selective assimilation (dissolution, ion exchange) without melting is difficult to ascertain and if complete is seen only as chemical evidence of contamination without any original physical evidence (Clarke, 2007; Vásquez et al., 2009). The above examples show the diversity of processes leading to different granitoid compositions by crystallization of mafic parents, partial melting of the crust, modification during migration, segregation and ascent, and different types of melt mixing and mingling. Granitoids are important probes to lithosphere and useful in crustal evolution studies but the complex origin has to be taken into account. The Paleoproterozoic Svecofennian orogeny (1.91–1.78 Ga) in Fennoscandia formed the composite Svecofennian orogen, which is one of the largest collage of Paleoproterozoic orogens globally (Lahtinen et al., 2005). Within the orogen, the Karelia province (Archean) is bounded across a cryptic suture zone (Koistinen, 1981), formed at 1.91–1.90 Ga, to the Svecofennia province (Paleoproterozoic) (Fig. 1). There is an abrupt change, based on the Sm–Nd and Pb–Pb data (Huhma, 1986; Lahtinen and Huhma, 1997; Vaasjoki, 1981), from Archean crust to juvenile arc rocks across the boundary. The ArcheanProterozoic (AP) boundary is characterized by anomalously thick crust (N 56 km), having a high velocity lower crust (Vp 6.8–7.3 km/s) between depths of 35–60 km (Korja et al., 1993), and thick layered lithospheric mantle (Lehtonen and O'Brien, 2009; O'Brien and Lehtonen, 2012; Peltonen and Brügmann, 2006). Mantle–crust decoupling of the Archean lithosphere in the AP boundary during collision is proposed with a formation of a crocodile jaw, where arc rocks are partly buried beneath the bounding Archean continent (see Lahtinen et al., 2015 and references therein). One noticeable feature in the central Fennoscandia is the Central Finland Granitoid Complex (CFGC in Fig. 1), which is a roundish area of approximately 40,000 km2 surrounded by supracrustal belts. Minor occurrences of gabbros and supracrustal rocks occur but the dominance of intermediate to felsic granodiorites and granites is a conspicuous feature. The granitoid rocks have typically U–Pb ages of 1.89–1.87 Ga (Huhma, 1986; Lahtinen and Huhma, 1997; Rämö et al., 2001) and have been further divided into syn-kinematic (1.89–1.88 Ga) and

post-kinematic (1.88–1.87 Ga) groups (Nironen et al., 2000; Rämö et al., 2001). The geochemical and isotope characteristics of the CFGC are interpreted to indicate an occurrence of an evolved 2.0 Ga lithosphere (Lahtinen, 1994; Lahtinen and Huhma, 1997; Rämö et al., 2001). This lithospheric block was named as ca. 2.0 Ga Keitele microcontinent by Lahtinen et al. (2005) but no direct evidence for a crust of this age has been observed so far (see Lahtinen et al., 2009). The opposite marginal basin hypothesis for the CFGC involves melting of large amounts of sedimentary rock and mafic oceanic crust with a possible contribution from Archean crust (Rutland et al., 2001, 2004; Williams et al., 2008). In Hietanen's (1975) plate tectonic interpretation for central Fennoscandia, the CFGC was an analog to the western Cordillera of North America. Currently continent-arc collision (ca. 1.91 Ga) between the Archean continent and the Paleoproterozoic microcontinent-arc collage is the favored model for the evolution of the AP boundary (Gaál, 1990; Kohonen, 1995; Lahtinen, 1994; Lahtinen et al., 2005, 2009, 2015; Nironen, 1997). Based on crustal scale seismic reflection surveys (Kukkonen et al., 2006), a crustal stacking model for the CFGC was proposed (Kukkonen et al., 2008; Sorjonen-Ward, 2006). Korja et al. (2009) interpreted, based on same seismic reflection data, that the CFGC is a deep, lower-level section of an old core complex (see also Korja and Heikkinen, 2008). In the results of analog modeling by Nikkilä et al. (2015) crustal stacking was followed by gravitational spreading via westward directed mid-crustal flow between 1.87 Ga and 1.86 Ga. Lahtinen et al. (2014) described two continuous large arcuate structures, a pair of coupled Bothnian Oroclines, in the central Fennoscandia (Fig. 1) coinciding with a crustal conductance anomaly (Korja et al., 2002), lithological continuations, and structural and metamorphic vergence. They suggested a buckling origin for the orocline formation, which would imply that the Keitele continental ribbon (CFGC) would have been originally a linear feature further buckled to its present equidimensional form. The deep seismic reflection profile FIRE 3A (Fig. 2; Kukkonen et al., 2006) crosscuts both the AP boundary and CFGC and images the present crustal structure. Along the profile, we combined geochemical data from samples collected for this study and samples used earlier for age dating with published geochemical data for plutonic rocks from the national rock geochemistry data base (Rasilainen et al., 2007). There are 12 key samples in which zircons were studied for inclusions, and analyzed (core-rim) by ion probe for U–Pb dating and oxygen isotopes. Later same zircons were analyzed for Lu–Hf with LA–MC–ICP–MS. TIMS and LA–MC–ICP–MS U–Pb data for zircons, monazites and titanites for 11 samples and new Sm–Nd data for 24 samples are also presented. Combined with previously published data, these will be used to assess the composition of crust and subcontinental lithospheric mantle (SCLM) in the Archean margin, AP boundary, CFGC and across the boundary zone to the Bothnia Belt (Figs. 1 and 2). In addition, we scrutinize the occurrence of possible sedimentary component in the granitoid magmas and the nature of their igneous source component, present new interpretation for the crustal profile along the FIRE 3A, and discuss about the possible tectonic implications based on our data and interpretations.

2. Geological setting and samples The study area (Fig. 2 and Appendix D) represents Archean basement and Paleoproterozoic cover rocks of the Karelia province and the Paleoproterozoic rocks of the Svecofennia province (Figs. 1 and 2). The study area along the FIRE 3A has been divided into the Archean continent (AC), Archean-Proterozoic boundary (AP), Central Finland Granitoid Complex (CFGC), boundary zone (BZ), and the Bothnia Belt (BB). Samples of this study have been divided into mafic rocks (gabbros, diorites, quartz diorites; SiO2 typically ≤60%), intermediate granitoids (SiO2 ≤ 68%) and felsic granitoids (SiO2 N 68%). Granitoids that have


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Fig. 1. A geological map of Fennoscandia modified after Koistinen et al. (2001). CFGC — Central Finland Granitoid Complex; BB — Bothnia Belt; PB — Pirkanmaa Belt; K — kimberlite field; SS — Southern Svecofennia; and CS — Central Svecofennia. Dashed green line — profile line for Fig. 12. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

Zr + Nb + Ce + Y N 400 ppm or Y + Nb ≥ 50 ppm (Pearce et al., 1984; Whalen et al., 1987) have been classified as A-type granitoids. The Paleoproterozoic sedimentary and volcanic cover successions of the Karelia province (AC) are termed Karelian formations (2.5–1.9 Ga; Laajoki, 2005). The rift-related sequences include shallow marine carbonates (2.1–2.05 Ga), mafic volcanic rocks and dykes, and pelites followed by turbidite deposits (Laajoki, 2005). The Siilinjärvi volcanic rocks extruded at ca. 2.06 Ga (Lukkarinen, 2008) and form the largest preserved rift to passive margin sequence in the study area. Shallow to deep water Karelian greywackes and mudstones (SA and AG in Appendix D), which are spread over on the western margin of the Karelia

Province are considered as Kalevian (2.06–1.87 Ga; Laajoki, 2005). The allochthonous (AG in Appendix D) deep water turbidites (Kontinen, 1987), with tectonically enclosed fragments of 1.95 Ga ophiolitic bodies (Peltonen, 2005), have variable maximum depositional ages from 1.95–1.94 Ga to 1.92 Ga (Lahtinen et al., 2010). The Paleoproterozoic granitoids in the AC occur as a 50–100 km wide zone east of the AP boundary, intrude both parautochononous and allochthonous sedimentary cover and Archean gneisses, and are typically ≤1.87 Ga in age (Huhma, 1986). The granitoids are often comingled with more mafic melts (pl-porphyrite RLL$-2009-1.1 and granodiorite 1.2 in Fig. 3) and are thus at least partly co-magmatic.

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Fig. 2. A simplified geological map of the study area modified from a Bedrock of Finland — DigiKP (see Appendix D for a more detailed map). Black lines — FIRE reflection seismic profiles. Red lines — block boundaries with CDP along FIRE 3A. Selected reference data from Rämö et al., 2001; Makkonen and Huhma, 2007; Lukkarinen, 2008; Nironen and Front, 1992; Ekdahl, 1993; Huhma, 1986. Black dots — granitoids and green dots — ultramafic-mafic plutonic rocks. Four digit numbers with errors define U–Pb zircon age in Ma and other numbers define εNd (1.88 Ga) values.

The felsic rocks are granodiorites and tonalites, and mafic-intermediate rocks occur as diorites and intermediate dykes. The latter have been referred to as microtonalites (Huhma, 1981) due to common occurrence of plagioclase phenocrysts. At least some of the large plutons have intruded accompanied by strike-slip deformation with overprinting structural strain (Gaál and Rauhamäki, 1971; Halden, 1982; Sorjonen-Ward, 2006), like Kermajärvi pluton (A25 in Fig. 2). Dextral displacement of at least 20 km is estimated for this deformation zone (Sorjonen-Ward, 2006). There is another transect of samples north of FIRE 3A, where many of the samples are from the Juurusvesi pluton (Fig. 2; Äikäs, 2000). Sorjonen-Ward (2006) interpreted this pluton as an example of lateral flow of magma from a nearby crustal-scale deformation zone. The Archean-Proterozoic (AP) boundary in Fig. 2 is gradational to both the AC and CFGC, where the western boundary towards the CFGC is defined by the occurrence of Older Svecofennia (1.93–1.92 Ga) juvenile island arc rocks (Kähkönen, 2005; Lahtinen, 1994; Lahtinen and Huhma, 1997; Vaasjoki, 1981). Both in the AP boundary and western part of the AC occur partly migmatitic paragneisses and amphibolites (MG in Appendix D), which have lately been included in the Karelia province (Lahtinen et al., 2015). The AP boundary partly coincides with the Raahe-Ladoga shear complex, a locally almost 100 km wide shear complex with dextral strike-slip tectonics (Kärki et al., 1993, 2012; Sorjonen-Ward, 2006). Deformation varies from early stage blastomylonites and migmatites to late-stage brittle structures. The complicated character of the AP boundary close to FIRE 3A profile is also seen in that the rocks of the Western Finland Supersuite, correlative to the Pirkanmaa belt in Fig. 1, and rocks of the Southern Svecofennia are emerging in the AP boundary in this area (Appendix D). Plutonic rocks in the AP boundary vary from gabbros to granodiorites, granites and quartz monzonites. Mafic enclaves and mingling between coeval felsic and mafic magmas are common and locally sedimentary enclaves are also found. Some A-type quartz monzonite/ granodiorites are porphyritic (A1969 in Fig. 3) and western intrusions

(e.g., A218) belong to Type 3 pyroxene-bearing post-kinematic plutons of Nironen et al. (2000). Deformation along the AP boundary (a shear complex) is strongly partitioned and the granitoids of same age vary from strongly deformed to non-foliated, e.g., A1968 and A1969 in Fig. 3, respectively. The Central Finland Granitoid Complex (CFGC) comprises Svecofennian supracrustal rocks (1.90–1.88 Ga) and plutonic rocks that mainly consist of foliated I-type (1.89–1.88 Ga) granodiorite and granite (Nironen, 2003; Nironen et al., 2000). Post-kinematic, undeformed or slightly foliated alkaline affinity plutons composed of quartz monzonites and monzogranites are common and have been divided into three types (Type 1 absent from the Fig. 2 study area; Elliott et al., 1998; Nironen et al., 2000; Rämö et al., 2001; Elliott, 2003). The two westernmost plutons of the porphyritic granitoids (Fig. 2) belong to the Type 2 and the plutons east of these to Type 3 in their classification. The A-type samples of this study along the FIRE 3A profile are normally from these porphyritic plutons (e.g., A1972; Fig. 2).Granodiorite A1971 and granite A545 are transitional between I- and A-types but are here classified as I-type granitoids. Granodiorite A1973 is even-grained and foliated (Fig. 3) example of the I-type granitoids. The boundary zone (BZ; Fig. 2) is characterized by abundant gabbros and that felsic granitoids are tonalitic to granodioritic in composition (Nironen, 2003). The BZ volcanic rocks have a calc-alkaline mature island arc affinity (Kähkönen, 2005). The structurally complicated and locally highly metamorphosed, dominantly sedimentary Bothnia Belt (BB, Figs. 1 and 2) includes a lower metabasalt sequence of WPB-MORB affinity, associated with metachert, black shales and metaturbidites, deposited at ~ 1.91–1.92 Ga (Chopin et al., 2014; Kähkönen, 2005; Vaarma and Pipping, 1997; Williams et al., 2008). The granitoids in the study area from the BB can be divided into older (≥1.90 Ga) and strongly foliated granitoids (e.g., A292 in Fig. 2), which either intrude the surrounding sediments or form a basement for them, and to crosscutting 1.89–1.88 Ga granitoids (A1153 in Fig. 3).


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Fig. 3. Selected sampled units. The magnet pen length is 13 cm and coin diameter is 2.3 cm. AC — Archean continent; AP — Archean-Proterozoic boundary; CFGC — Central Finland granitoid complex; BZ — boundary zone. /M — mafic rock; /F– felsic granitoids; /G — intermediate granitoids; and /A — A-type granitoids.

The latter ones can be correlated with the BZ granitoids. The core of the BB (outside of Fig. 2) is composed of the Vaasa complex having in-situ sedimentary derived 1.88–1.87 Ga granites that are interpreted to present a magmatic dome (Kotilainen et al., 2016). Youngest plutonic rocks in the BB are the ca. 1.80 Ga pegmatite granites, which are left out of discussion. 3. Analytical techniques Major elements and Ba, Cr, Cu, Ni, S, Sr, V, Y, Zn, and Zr were determined by XRF from pressed briquettes, and non carbonate C and CO2 by Leco CR-12 carbon analyzer. REE, Co, Nb, Hf, Rb, Sc, Ta, Th and U were determined by ICP–MS from solution combining those after dissolution of the sample (0.2 g) with hydrofluoric acid-perchloric acid and a lithium metaborate/sodium perborate fusion of the insoluble residue. Two internal quality control samples (granite QC-1 and peridotite QC-2) were analyzed for each analysis batch. In addition to these control samples, certified reference materials and in-house reference samples were analyzed periodically to provide information on the quality of the measurement data (see Rasilainen et al., 2007). Analytical techniques for Sm–Nd data have been described in Lahtinen and Huhma (1997). The long-term average 143Nd/144Nd for the La Jolla standard is 0.511851 ± 0.000008 (standard deviation for 69 measurements during 2012) and based on duplicate analyses, the error in 147Sm/144Nd is estimated to be 0.4%. U–Th–Pb analyses of zircons in key samples performed in 2011 using the CAMECA IMS

1280 ion-microprobe (NORDSIM) located at the Swedish Museum of Natural History, Stockholm (Whitehouse and Kamber, 2005; Whitehouse et al., 1999). Calibration of the U/Pb ratio was based on analyses of the Geostandards zircon 91500, which has an age of 1065.4 ± 0.3 Ma (Wiedenbeck et al., 1995). Corrections for common Pb are based upon the measured 204Pb signal, and the present day terrestrial average Pb-isotopic composition (Stacey and Kramers, 1975). NORDSIM data reduction employed Excel routines developed by Whitehouse and program by Andersen was used for ICPMS data (Andersen et al., 2004). Analyses of δ18O were performed using the same instrument in close proximity to the dated spots after re-polishing. The instrument setup and analytical procedures for oxygen isotopes in zircon followed closely those of Whitehouse and Nemchin (2009), utilizing a ca. 2 nA Cs + primary ion beam together with a normal incidence low energy electron gun for charge compensation and two Faraday detectors at a mass resolution (M/ΔM) of ~2500. Measurements were performed in pre-programmed chain analysis mode with automatic field aperture centering on the 16O signal. The magnetic field was locked using NMR regulation. Each data-acquisition run comprised a 20 × 20 μm presputter to remove the Au layer followed by beam centering and 64 s of data integration performed using a non-rastered, ca, 10 × 10 μm spot. Field aperture centering values were well within those for which no bias has been observed during tests on standard mounts (Whitehouse and Nemchin, 2009). A total of 92 unknowns were measured during two analysis sessions with every set of six unknowns

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bracketed by two analyses on the Geostandards 91500 zircon. A δ18O value of +9.86‰ (SMOW, Wiedenbeck et al., 2004) was assumed for the 91500 zircon and a small drift corrections was applied where necessary, based on the regular standard measurements. External reproducibility of ±0.21‰ (1 SD) based on the standard measurements was propagated onto the overall uncertainty for each analysis reported in Appendix B. Some of the U–Pb analyses were performed at GTK using a Nu Plasma HR multicollector ICP–MS combined with a New Wave UP193 or Photon Machine Analyte G2 (for A2100) laser microprobe. Standard zircons GJ-1 (609 ± 1 Ma) and in-house standards A382 (1977 ± 2 Ma) or A1772 (2712 ± 1 Ma) were used for calibration (Huhma et al., 2011, 2012). Lu–Hf isotope analyses were performed using the same instruments. All analyses were made using the following parameters; beam diameter: 40 to 50 μm, pulse frequency: 5 Hz, beam energy density: 2.8 J/cm2. Each ablation was preceded by a 30 s on-mass background measurement. During the ablation the following masses were collected in static mode: 171Yb, 172Yb, 173Yb, 175Lu, 176Hf–Yb–Lu, 177Hf, 178 Hf, 179Hf. The total Hf signal obtained for zircons was typically 1.0–2.0 V. Isotopic ratios were measured using the Nu Plasma timeresolved analysis software. The isotopic ratios were later on calculated off-line using an Excel spreadsheet. The raw data were filtered at 2σ and corrected for mass discrimination using an exponential law. The mass discrimination factor for Hf was determined assuming 179 Hf/177Hf = 0.7325 (Patchett et al., 1981). The mass discrimination factor for Yb was determined assuming 173Yb/171Yb = 1.132685 (Chu et al., 2002). The 176Lu/175Lu value of 0.02656 has also been used for the correction of the 176Lu interference on 176Hf (Scherer et al., 2001; Vervoort et al., 2004). A value for the decay constant of 176Lu of 1.867 × 10− 11 a− 1 has been used in all calculations (Scherer et al., 2001; Söderlund et al., 2004). For the calculation of εHf values we use present-day chondritic 176Hf/177Hf = 0.282785 and 176Lu/177Hf = 0.0336 (Bouvier et al., 2008). The zircon standard GJ-1 was run at frequent intervals for quality control. Multiple LA–MC–ICP–MS analyses, using the same instrumental parameters, of the reference zircon GJ-1 during the course of the present study yielded a 176Hf/177Hf of 0.28203 ± 5 (1 s, n = 16, which is just within error to results obtained by solution MC–ICP–MS analyses for GJ1 (0.281998 ± 7, Gerdes and Zeh, 2006; 0.282000 ± 5, Morel et al., 2008). The U–Pb TIMS analyses followed standard procedures (Huhma et al., 2012). Age calculations and related diagrams were made using Isoplot/Ex v 2.49 (Ludwig, 2003). 4. Results All the geochemical data are presented in Appendix A, SIMS data (U–Pb and O) in Appendix B, and LA–MC–ICP–MS (U–Pb, Lu–Hf) and TIMS data (U–Pb) in Appendix C, respectively. 4.1. Zircon morphology and inclusions within zircons Table 1 shows list of samples used for the SIMS and Lu–Hf and O isotope studies which samples were also studied in detail for inclusions in zircon. Tonalite A1756 was included in the inclusion studies. Granitoids from the AC have typically light-colored, euhedral and zoned zircons. Zircons are often elongated (e.g., A24 in Fig. 4) and have average length/ width ratio (l:w) of 4–6:1. Needle- and crystal-shaped apatites are main phases of inclusions in zircons of granodiorite A24 (Fig. 4) and tonalite A1756, whereas quartz and orthoclase inclusions characterize granodiorites A25 and A60 (Fig. 4). Granite A218 and tonalite A1968 zircons from the AP boundary show oscillatory zoning with l:w ratio 1.2–2, and have apatite as the main inclusion. Granite A218 includes feldspars and tonalite A1968 amphiboles as minor inclusion phases. The main inclusion phase in granodiorite A1969 zircons is biotite (Fig. 4). Zircons in granitoids of the CFGC show typically oscillatory zoning with l:w ratio 1.2–2, and have apatite as predominant inclusion phase.

Orthoclase, quartz and minor occurrences of biotite and amphibole are also locally present. Main exception is granite A545, where the main phases of zircon inclusions are compound minerals and undefined feldspar-Fe-mineral-mix, interpreted as hematite micro-inclusions within the feldspar (Fig. 4). Quartz, feldspars, biotite, apatite, hematite and ilmenite occur as crystals or in compounds. Monazites (Fig. 4) and xenotime occur as ovoids and crusts around inclusion nodules. Granodiorite A292 from the BB has ovoid, needle- and crystal-shaped apatites as main phases of zircon inclusions with titanite, feldspar, quartz and biotite as minor inclusions. 4.2. Age data All the age data are presented in Appendixes B and C and summarized in Table 1. Concordia diagrams for samples analyzed by ion probe are shown in Fig. 5. One possible older core of ca. 1914 ± 8 Ma was found from the sample A1970. Granodiorite A1973 shows scatter in zircon ages from 1875 ± 9 Ma to 1904 ± 10 (Fig. 5; Appendix B), and a single xenocrystic Archean zircon was found from each sample A25 and A1968. Four granitoids from the AC (A23, A24, A25, A1756) have precise zircon U–Pb ages between 1870 and 1865 Ma, whereas granodiorite A60 and plagioclase-porphyritic dyke A1574 have younger zircon U–Pb ages of 1850 ± 2 Ma and 1833 ± 7 Ma, respectively. Monazite and titanite U–Pb ages are normally 5–10 Ma younger than obtained zircon ages. The ion probe U–Pb data for three samples (A218, A1968, A1969) from the AP boundary are between 1879 and 1882 Ma, and sample A146 (1887 ± 6 Ma) is within error limits comparable with them. The CFGC granitoids show an ion probe age variation from 1876 ± 6 Ma (A1972) to 1891 ± 5 Ma (A545). Granodiorite A447b has an age of 1875 ± 3 Ma comparable to A1972, whereas other granitoids, excluding A545, have ages between 1883 and 1887 Ma. Three boundary zone samples (A380, A579, A926) have a tight cluster of ages at 1881–1883 Ma. Granodiorite A1153 (1891 ± 15 Ma) from the BB is included as a boundary zone sample as it seems to be younger, although having a large error than the other dated granodiorites A1270b (1898 ± 2 Ma) and A292 (1921 ± 5 Ma) in the BB. 4.3. Sm–Nd and Lu–Hf data The Sm–Nd data are shown in Table 1 and Lu–Hf data in Appendix C. Samples from the AC show variable negative epsilon (T) values with granodiorite A60 having the most negative εNd (1.85 Ga) of − 6.1 (Fig. 6a). The AP boundary granitoids have typically slightly negative epsilon values, but tonalite A1968 shows a positive εNd (1.88 Ga) of +1.2. The CFGC granitoids εNd (T) values vary between −1.7 and 1.1 with most values around 0. The BZ εNd (T) values are between + 1.8 and + 2.9 with lower value of + 0.5 in tonalite A1153. The two older granitoids of the BB have values along the DM line (Fig. 6a). The average hafnium isotope data from zircons correlate well with whole rock Sm − Nd data (Fig. 6b) but there is a large variation in the zircon epsilon-Hf values of the AC granitoids (A24, A25 and A60). This internal variation within the samples is correlated with a decrease in εNd value, which indicates that an increasing percentage of recycling component could be responsible for that large variation in εHf. 4.4. Oxygen isotope data The oxygen isotope data are presented in Appendix B and results are also shown in Fig. 7, where the MORB range and “normal δ18O” magma differentiation array in igneous rocks from Bindeman et al. (2004) are also presented (see also Bindeman, 2008). The zircons from the AC samples (A24, A25, A60) were long and fractured and thus corerim pairs are lacking. The data show mainly values in the Bindeman range with few elevated values for the three AC granitoids. The AP granitoids (A218, A1968, A1969) show higher values and, especially,


513

Table 1 Sm-Nd isotope data for the plutonic rocks of this study. Samples analyzed by SIMS and used for Lu–Hf and O studies are marked bold. Areaa

Sample

A23 A24b A25 A60 A82 A85 A105 A146 A172 A218 A292 A380 A447b A545 A579 A851 A926a A940 A1153 A1270b A1574 A1614 A1756 A1968 A1969 A1970 A1971 A1972 A1973 RLL$-2009–1.1 RLL$-2009–1.2 RLL$-2009–3 RLL$-2009–12 RLL$-2009–19

Granodiorite Granite Granodiorite Granodiorite Granite Granodiorite Quartz diorite Quartz diorite Granodiorite Granite Granodiorite Gabro Granodiorite Granite Granodiorite pl-porphyrite dyke Gabbro pl-porphyrite dyke Tonalite Granodiorite pl-porphyrite dyke Syenite Tonalite Tonalite Granodiorite Granodiorite Granodiorite Granite Granodiorite pl-porphyrite Granodiorite Diorite Quartz monzonite Granodiorite

AC/G AC/F AC/F AC/F CFGC/G CFGC/A AP/G AP/G AC/M AP/A BB/F BZ/M CFGC/G CFGC/F BZ/F AC/M BZ/M AC/M BZ/G BB/F AC/M AC/G AC/G AP/G AP/A CFGC/G CFGC/F CFGC/A CFGC/F AC/M AC/F AC/M CFGC/A BZ/G

Finnish grid ccoordinates

Sm ppm Nd ppm

x

y

6,999,000 6,929,300 6,936,800 6,991,573 6,939,941 6,939,804 6,983,980 6,980,030 6,990,500 6,935,140 7,044,269 7,111,042 6,906,935 6,915,270 7,101,760 6,993,970 7,056,606 6,997,110 7,033,647 7,102,912 6,992,720 7,003,710 6,999,582 6,924,500 6,921,650 6,936,600 6,958,250 6,977,580 7,005,156 6,935,014 6,935,014 6,928,878 6,941,018 7,022,235

3,548,000 2.83 3,559,800 3.66 3,577,000 5.12 3,586,107 5.88 3,355,197 4.92 3,364,603 8.50 3,511,350 4.42 3,500,530 4.61 3,516,860 4.11 3,501,730 8.73 3,339,440 4.07 3,375,373 2.05 3,383,302 3.72 3,435,730 7.82 3,359,225 2.54 3,545,800 9.86 3,363,151 3.50 3,586,760 7.80 3,320,356 2.17 3,341,585 2.48 3,544,410 11.60 3,540,680 2.72 3,555,093 2.62 3,523,720 2.96 3,523,820 9.03 3,472,700 4.62 3,409,310 5.08 3,391,840 7.72 3,372,406 3.05 3,604,538 10.44 3,604,538 4.83 3,594,598 7.83 3,437,447 9.19 3,363,130 2.09

147 144

18.99 22.75 40.44 39.62 27.82 55.40 25.63 25.34 24.96 43.15 22.34 6.74 19.00 32.91 17.09 61.39 15.30 52.18 11.28 14.82 71.80 15.00 13.19 17.36 51.81 25.09 23.66 36.91 18.40 74.89 37.14 45.94 44.23 12.99

Sm/ Nd

0.0901 0.0971 0.0767 0.0896 0.1068 0.0928 0.1043 0.1099 0.0999 0.1223 0.1100 0.1844 0.1182 0.1436 0.0899 0.0971 0.1386 0.0904 0.1167 0.1012 0.0976 0.1096 0.1203 0.1033 0.1054 0.1114 0.1299 0.1264 0.1003 0.0842 0.0785 0.1029 0.1256 0.0974

143 144

Nd/ Nd

0.511171 0.511243 0.511056 0.511023 0.511559 0.511412 0.511328 0.511479 0.511327 0.511690 0.511738 0.512617 0.511658 0.511934 0.511468 0.511384 0.512006 0.511261 0.511669 0.511632 0.511348 0.511383 0.511537 0.511546 0.511457 0.511551 0.511728 0.511754 0.511447 0.511224 0.511033 0.511331 0.511776 0.511504

2 sigma error eNd(T) TDM Ma U–Pb ageb Ma zircon 0.000010 0.000034 0.000022 0.000035 0.000018 0.000021 0.000026 0.000022 0.000010 0.000010 0.000010 0.000017 0.000020 0.000034 0.000033 0.000010 0.000010 0.000010 0.000010 0.000011 0.000010 0.000010 0.000010 0.000010 0.000010 0.000010 0.000010 0.000010 0.000011 0.000010 0.000010 0.000010 0.000010 0.000010

-3.1 -3.4 -2.1 -6.1 0.7 1.1 -3.3 -1.7 -2.3 -0.6 3.8 2.5 -0.2 -0.9 2.9 -0.6 1.6 -1.4 0.5 3.7 -1.9 -3.4 -3.2 1.2 -1.1 -0.7 -1.7 -0.4 0.0 -0.6 -3.0 -3.1 0.3 1.8

2304 2353 2208 2483 2114 2052 2391 2297 2292 2254 1913 2210 2421 1935 2161 2099 2196 2156 1906 2219 2427 2456 2062 2229 2219 2393 2249 2141 2137 2262 2355 2192 2013

U–Pb ageb Ma mz/tit

Ref.c

1870 ± 2 a 1868 ± 7 1863 ± 3 a, 1 1865 ± 6 a, 1 1850 ± 6 1843 ± 3 a, 1 1886 ± 10 1880 ± 10 1 1883 ± 17 1 ~1870 1850 ± 7 4 1887 ± 6 1852 ± 7 4 ~1870 1869 ± 19 4 1879 ± 2 a, 2 1921 ± 5 a, 3 1883 ± 8 5 1875 ± 3 a, 1 1891 ± 5 a, 1 1883 ± 5 1 ~1870 a 1881 ± 6 6 ~1870 1850 ± 7 7 1891 ± 15 8 1898 ± 2 a 1833 ± 7 1829 ± 13 a 1891 ± 23 1859 ± 23 a 1869 ± 5 1857 ± 2 a 1882 ± 6 a 1882 ± 5 a 1883 ± 6 a 1885 ± 3 a 1876 ± 6 a 1887 ± 4 a ~1870 a ~1870 a ~1870 a ~1880 a ~1880 a

TDM calculated according to DePaolo (1981). a AC — Archean continent; AP — Archean-Proterozoic boundary; BZ — Boundary zone; CFGC — Central Finland Granitoid Complex; BB — Bothnian Belt; /M — mafic; /G — intermediate; /F — felsic; and /A — A-type. b Zircon age data by SIMS (this study, bold); other zircon, monazite and titanite ages by TIMS, SIMS or LA–MCICPMS; ages with ~ are inferred ages. Mz — monazite; and tit — titanite. c Refs. a — this study; 1 — Huhma, 1986; 2 — Korsman et al., 1984; 3 — Lahtinen and Huhma, 1997; 4 — Lukkarinen, 2008; 5 — Patchett and Kouvo, 1986; 6 — Kärkkäinen, 1999; 7 — Huhma, 1981; and 8 — Vaasjoki et al., 2003.

the granodiorite A1969 values are clearly outside the Bindeman range. Noticeable is that the measured core-rim pairs indicate lower δ18O values for A1968 rims and higher values for A218 and A1969 rims. The granitoids A1970, A1971 and A1972 in the CFGC have values within or close to the Bindeman range whereas granite A545 has slightly higher values. Granodiorite A1973 has the biggest core-rim differences, where most of the cores are slightly above the Bindeman range with rims having substantially lower values. The granodiorite A292 from the BB shows low δ18O values plotting in the MORB range (Fig. 7). 4.5. Geochemistry Raw geochemical data are presented in supplementary Appendix A, and sample points and selected Harker diagrams in Appendix D. Some geochemical parameters along the FIRE 3A profile are shown in Fig. 8 and spidergrams for selected samples in Fig. 10. The studied rock samples have been separated to mafic, intermediate and felsic rocks and a geochemical discrimination (Zr + Nb + Ce + Y N 400 ppm or Y + Nb ≥ 50 ppm; Pearce et al., 1984; Whalen et al., 1987) has been used to separate A-type granitoids from I-type granitoids to own intermediate and felsic groups. These A-type rocks have, higher Zr, Nb and Y, higher FeOT and TiO2 and Zn, and lower MgO, CaO and Sr showing thus the typical A-type/within-plate granitoid signatures (e.g. Pearce et al., 1984; Whalen et al., 1987; Kilpatrick and Ellis, 1992; Appendixes A and D).

The aluminium saturation index (ASI, Fig. 9) increases towards higher SiO2 values and most samples from the AC, CFGC, BZ and BB follow the predicted igneous trend. On the contrary, the A-type granitoids from the AP boundary follow the sediment assimilation trend of mixing between mafic magma, and sedimentary material and/or derived melt. Also few other granitoids from the AP boundary follow this trend. The plutonic rocks in the AC show a bimodal nature where the more mafic rocks are typically intermediate in composition (Fig. 8) and high-K and tholeiitic, whereas the granitoids are medium-K to high-K and calc-alkaline (Appendix D). Other characteristic features are that they are typically both enriched in Ba and Sr (Appendix D; Fig. 10) and have high K/Rb and Sr/Y (Fig. 8). Typically they also have high normalized La/Yb and Gd/Yb ratios, elevated Th/U and some have positive Eu anomalies (Appendix D). The AP boundary is characterized by mafic rocks and intermediate granitoids (Fig. 8). Mafic rocks are often gabbros and ultramafic tholeiitic intrusions with SiO2 b 55% and two of the more intermediate mafic rocks show very high ASI values relative to their SiO2 (Fig. 9). The I-type granitoids are heterogeneous but typically medium-K calc-alkaline granitoids, and characterized by high Sr/Y (Fig. 8) and high normalized La/Yb and Gd/Yb ratios (Appendix D). Some plot along the sediment assimilation trend in Fig. 9 and have high Th/U ratios (5.7–7.6) and some (e.g., A1968 in Fig. 10) show elevated Nb/Ta (Fig. 8) and low Th/U (1.1– 2.9). A-type granitoids are high-K to very high-K tholeiitic rocks and mainly intermediate in composition (Fig. 8; Appendix D). They follow

514


Fig. 4. SEM-images of selected zircons with inclusions. See Fig. 3 for legend.

the sediment assimilation trend (Fig. 9), have low Sr/Y ratios (Fig. 8), Ba minimum (Fig. 10), and typically high Th/U ratios (3–12; Appendixes A and D).

The CFGC plutonic rocks are dominated by felsic granitoids, whereas mafic rocks with basaltic composition (SiO2 b 52%) are rare (Fig. 8). Both mafic and felsic rocks have typically low Nb/Ta ratios (Fig. 8) in the

Fig. 5. Concordia diagrams for SIMS analysis of key samples (bold in Table 1). AC — Archean continent; AP — Archean-Proterozoic boundary; CFGC — Central Finland granitoid complex; BB — Bothnian belt. /F — felsic granitoids; /G — intermediate granitoids; and /A — A-type granitoids.


515

Fig. 6. a. Sm–Nd isotope diagram showing data from samples in this study (Table 1). b. Comparison of whole rock Sm–Nd data and Lu–Hf data in zircons from key samples (bold in Table 1). Green lines with sample numbers and Th/Yb ratios in 6a are data for the 2.06 Ga Siilinjärvi basalts (Lahtinen et al., 2015). 1.1 and 1.2 refer to RLL$-2009-1.1 and 1.2, respectively. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

central part of the CFGC, which low ratio correlates with high Th (see central CFGC samples A1971 and A1972 in Fig. 10). A-type granitoids are typically very high-K and tholeiitic with high K/Rb ratios in the intermediate granitoids (Fig. 8; Appendix D). The felsic A-type granitoids characterize the western part of the CFGC (Fig. 8). The I-type granitoids are high-K and the intermediate granitoids are calc-alkaline (Appendix D). The felsic (SiO2 N 68%) I-type granodiorites and granites show increase in FeO/MgO ratio and strong decrease in Ba and Sr contents and increase in negative Eu anomaly in Harker diagrams (Appendix D) which indicates differentiation with strong feldspar fractionation (e.g., A1971 in Fig. 10). The felsic A-type granitoids have similar but less steep decreasing trend in Sr and increase negative Eu anomaly but Ba typically remains more constant (Appendix D). The strong Sr fractionation explains the differences in Sr/Y ratios (Fig. 8) where the mafic rocks and intermediate I-type granitoids show high ratios but felsic I-type granitoids low ratios. The boundary zone (BZ) has a bimodal distribution of mafic plutons and more felsic granitoids and this bimodality is seen also in the Nb/Ta ratio (Fig. 8). Granitoids are typically calc-alkaline and I-type, and compared to the CFGC mainly medium-K with also lower Rb–Ba–Th values (e.g., RLL$-2009-19 Fig. 10) and higher K/Rb (Fig. 8) than in the adjacent CFGC granitoids. The two intermediate granitoids in the BB (Fig. 8) are

from tonalite pluton A1153 included with BZ granitoids. The other granitoids from the BB are also calc-alkaline and characterized by high Na2O, low K2O and low Rb (Appendix D), and also high K/Rb ratio (Fig. 8). The three high Sr/Y samples (Figs. 8), from a tonalite pluton (A1270; Table 1 and Fig. 10) north of the FIRE 3A profile (Fig. 2), have higher Al2O3, CaO and Sr level in Harker diagrams, and high normalized La/Yb and Gd/Yb values (Appendix D). Two low Sr/Y samples (Fig. 8) are from the oldest granitoid (1.92 Ga) in this study (A292 in Fig. 10) and differ from the other I-type granitoids having also low normalized La/Yb and Gd/Yb values (Appendix D). 5. Discussion 5.1. Mafic rocks and mantle composition Granitoids were the main focus of this study, without targeting the mafic suites. Most of the mafic samples available for this study are fractionated and intermediate in composition. The Th/Yb-Ta/Yb diagram (Fig. 11) shows that some of the more mafic rocks (SiO2 b 52%) from the AP boundary plot in the mantle array and also that the mafic rocks are dominantly from an enriched mantle source. The deviation to higher Th/Yb values in Fig. 11 can be due to an addition of a subduction

Fig. 7. Oxygen isotope SIMS data in zircons from key samples (bold in Table 1). MORB range and “normal δ18O” magma differentiation array from Bindeman et al. (2004). AC — Archean continent; AP — Archean-Proterozoic boundary; CFGC — Central Finland Granitoid Complex; and BB — Bothnian Belt.

516


Fig. 8. Selected geochemical parameters of samples in this study projected to the FIRE 3 A seismic reflection line (Fig. 2). The projected CDP points for individual samples are shown in Appendix A. M — mafic rocks; G — intermediate I-type granitoids ( SiO2 ≤ 68%); F — felsic I-type granitoids ( SiO2 N 68%); A — intermediate A-type granitoids (SiO2 ≤ 68%); and AF — felsic A-type granitoids ( SiO2 N 68%).

component to mantle-derived magma, melting of mantle contaminated earlier by a subduction process or contamination by continental crust. Schematic cross-section of the lithosphere across the AP boundary (Fig. 12) is based on geophysical data (Hyvönen et al., 2007; Yliniemi et al., 2004) and xenocryst data from 600 to 1200 Ma kimberlites (Lehtonen and O'Brien, 2009; O'Brien and Lehtonen, 2012; Peltonen and Brügmann, 2006). The mafic rocks of the AC are mainly intermediate in composition and enriched in Ba, Sr, LREE, Ti and P (Fig. 10). The

mafic rocks show also a Nb–Ta minimum but e.g. in sample A851 (Fig. 10) this minimum is not very pronounced. The εNd (T) values are between −0.6 and −1.9 (Table 1) and thus, clearly below DM values (Fig. 6). Close to the 600 Ma kimberlites (western K in Fig. 1) occur 1.79 Ga shoshonitic lamprophyre dykes, which are also enriched in Ba, Sr, LREE, Ti and P with clear negative Nb–Ta anomaly (Woodard et al., 2014). Metasomatic enrichment of mantle lithosphere by alkaline silicate melts followed by carbonatitic melts, both deriving from


Fig. 9. Aluminum saturation index diagram for plutonic rocks from this study showing also inferred igneous and sediment assimilation trends. A/CNK is the molar ratio Al2O3 / (CaO + K2O + Na2O). AC — Archean continent; AP — Archean-Proterozoic boundary; AP/A — A-type granitoids from the AP boundary; CFGC — Central Finland granitoid complex; CFGC/A — A-type granitoids from the CFGC; BZ — boundary zone (BZ); and BB — Bothnian Belt.

subducting sediments, have been proposed for the origin of the enriched mantle from where the lamprophyres derive (Andersson et al., 2006; Woodard and Huhma, 2015; Woodard et al., 2014). Similar origin could be plausible for the 1.87–1.83 Ga mafic-intermediate rocks in the AC, but the AC mafic rocks are high-K in character but not

517

Fig. 11. Th/Yb — Ta/Yb diagram (Pearce, 1983) for mafic rocks in this study. Siilinjärvi basalts after Lahtinen et al. (2015). AC — Archean continent; AP — Archean-Proterozoic boundary; CFGC — Central Finland Granitoid Complex; and BZ — boundary zone.

shoshonitic. Another possibility is a contamination of mantle melts with felsic melts via selective assimilation (dissolution, ion exchange; Clarke, 2007; Vásquez et al., 2009) to produce the observed characteristics, e.g., compositional similarity with the AC felsic rocks (Fig. 10) and decreasing εNd (T). The Siilinjärvi 2.06 Ga basalts witness a crustal contamination (Figs. 6 and 11) with Archean crust producing basalts that have very negative εNd values and have “subduction”-type Nb–Ta minimum in spidergrams (Fig. 6 in Lahtinen et al., 2015). This indicates

Fig. 10. MORB-normalized (Pearce, 1983) element patterns for selected samples of this study. M — mafic plutonic rocks; F — felsic granitoids; G — intermediate granitoids; and A — A-type granitoids.

518


Fig. 12. Schematic cross-section of the lithosphere across the AP boundary (Fig. 1) based on geophysical data (Hyvönen et al., 2007; Kozlovskaya et al., 2004; Yliniemi et al., 2004) and xenocryst data from 600 to 1200 Ma kimberlites (Lehtonen and O'Brien, 2009; O'Brien and Lehtonen, 2012; Peltonen and Brügmann, 2006). K — kimberlite fields in Fig. 1. SCLM — Sub-continental lithospheric mantle.

how difficult it is to separate the subduction component fingerprints from crustal contamination. The proposed mantle source for the AC mantle melts is the 2.1–2.0 Ga metasomatized lower part of the Archean SCLM (Fig. 12) with further contamination of these melts by continental crust before entering the middle-upper crust. A derivation from 1.90– 1.88 Ga metasomatised Paleoproterozoic SCLM cannot be ruled out. According to Makkonen and Huhma (2007) mafic-ultramafic intrusions along the eastern part of the AP boundary show a range in initial εNd values at 1880 Ma from − 2.4 to + 0.4 and western part from + 1.1 to + 1.7 and one occurrence in the CFGC from + 1.7 to + 1.9 (see also Fig. 2). They interpreted this variation in initial εNd values as contamination of EMORB parental magmas with Archean material either directly or via sedimentary material. Two gabbros from the BZ have εNd (T) values of +1.6 and +2.5 (Table 1) and have compositions of high-Ti tholeiitic basalt showing a rather juvenile character. Some gabbros and diorites in the AP are also very Ti–P rich (Appendix A) and have compositions of mildly alkaline basalt and potassic trachybasalt similar to the AC mafic rocks. Based on Fig. 12, it is possible that the mafic rocks from the eastern part of the AP boundary derive from the 2.1–2.0 Ga metasomatized SCLM and mafic rocks more to the west are either from the same metasomatized mantle and/or from the Paleoproterozoic SCLM. Unfortunately, kimberlites have not yet been found from the Paleoproterozoic part inhibiting direct sampling of the mantle. 5.2. Sedimentary input The AC granitoids often intrude sedimentary rocks and have locally sedimentary enclaves. The variation towards few higher δ18O values above the Bindeman range (Fig. 7), composed of MORB range and “normal δ18O” magma differentiation array in igneous rocks, could be interpreted as evidence for an involvement of sedimentary component. On the other hand, most of the AC granitoids follow the igneous trend (Fig. 9) and have strong positive Ba anomaly (Fig. 10) which contradicts with the strong negative Ba anomaly found in sedimentary rocks and igneous rocks assimilated with sedimentary material (Lahtinen and Huhma, 1997). It seems that sediment assimilation has occurred locally but the composition of the AC granitoids is considered to reflect mainly their igneous sources. The situation is totally different with the AP A-type granitoids (A218 and A1969), which plot towards the sediment assimilation trend (Fig. 9), have high δ18O values (6.23–8.25‰; Fig. 7) and exhibit negative Ba anomalies (Fig. 10). These all favor a large sedimentary component in their genesis. Elliott et al. (2005) have analyzed oxygen isotopes from

zircons in A218 (Fig. 2) by laser fluorination method and their δ18O (Zrc) of 7.71 is similar to the three highest values in this study (7.39– 7.68 ‰; Fig. 7). The few core-rim pairs in zircons of the A-type granitoids A218 and A1969 are overlapping within error but the rims have on average slightly higher δ18O values than the cores. This could indicate at least two stage sedimentary inputs, where melts having already strong sedimentary component further assimilate sedimentary material. The granitoids A218 and A1969 show negative εNd (T) values of − 0.6 and − 1.1 and TDM values of 2229 and 2254 Ma, respectively (Table 1). The metasedimentary rocks at the present crustal level in the AP (MG in Appendix D) show TDM values between 2.4 and 2.5 Ga (Lahtinen et al., 2015) and could present the sedimentary component deeper in the crust. Granitoid A218 and surrounding A-type granitoids have been classified by Nironen et al. (2000) as C-type intrusions with partial melting of a mafic granulitic residue to produce these magmas (Kilpatrick and Ellis, 1992). The granitoids have typically a pyroxene ± olivine-bearing marginal assemblage, which indicates relative reducing conditions and pressures on the order of 5–7 kbar during crystallization (Elliott et al., 1998). The AP A-type granitoids are typically intermediate in composition (Figs. 8 and 9) and a purely crustal-derived melt origin is not considered viable (cf. Patiño Douce, 1999). We propose that mantlederived melts caused dehydration melting of granulite-facies sedimentary migmatites and mixed or assimilated with these melts in mid-crustal magma chambers (depth 18–25 km; Elliott et al., 1998). Open-system fractional crystallization from a mafic magma with associated selective assimilation from sedimentary material has been proposed for the generation of some fayalite-bearing granitoids (Vásquez et al., 2009), which explains also many features noticed here. During ascent further assimilation of sedimentary material occurred, seen in higher δ18O values in the zircon rims (this study), and fractionation and progressively increasing oxidation state of the evolving, residual magma from which the central parts of the pluton eventually crystallized (Elliott et al., 1998). The late amphibole crystallization, presumably at 10–15 km, could register conditions at the level of emplacement (ibid). Some mafic rocks and also few I-type granitoids from the AP either follow the sediment assimilation trend or deviate towards it in Fig. 9. Apart few outliers, the CFGC, BZ and BB mafic rocks and granitoids follow the igneous trend. Granitoids A1970–A1972 from the CFGC and A292 from the BB have δ18O values (4.76–6.49‰) suggesting purely mantle-igneous source (Fig. 7). Granodiorite A1973 zircon cores have slightly elevated δ18O values, which could indicate minor sedimentary input. Granite A545 has also slightly elevated δ18O values and hematite micro-inclusions in zircons with monazites and xenotime locally as


ovoids and crusts around inclusion nodules (Fig. 4). This could indicate that the source material has been surficially altered (oxidized) igneous granitoid (feldspar inclusions) or derived oxidized arkose source. 5.3. Igneous source composition The in general bimodal nature of the AC magmatism (Fig. 8) and differences in the Sm–Nd characteristics of mafic and felsic melts (Fig. 6) exclude the possibility of felsic magmas to be solely a result of fractionation of mantle-derived melts. There is a large variation in the Ba/Rb and Ba/Th ratios, Eu anomaly and Sm–Nd (Appendix D and Figs. 6 and 10) favoring variable source components for the felsic melts and possibly cumulate origin for some tonalites (very positive Eu anomaly). Anyhow, granitoids with high K/Rb, Ba/Rb and Ba/Th (like RLL$-2009-1.2 and A60 in Fig. 10) should have a source with similar ratios, as these elements are normally incompatible with melt residue. Both lower continental crust (Rudnick and Gao, 2003) and enriched mantle-derived melts (Pearce, 1983) show the above-mentioned characteristics. In the Zr/Hf–Nb/Ta diagram (Fig. 13) many of the AC granitoids plot above the field of continental crust. The average lower crustal values of 35.8 (Zr/Hf) and 8.3 (Nb/Ta) calculated from Rudnick and Gao (2003) plot lower crust below the field of continental crust in Fig. 13. John et al. (2011; see also Münker et al., 2003) proposed that a hydrous eclogite or lower crustal high-pressure amphibolites produce high Nb/Ta melts at low degrees of melting at depths between 30 and 70 km, where titanite and rutile can coexist. Thus, one possible melt source would be a low degree melting of eclogites or garnet-bearing amphibolites from an enriched mantle source (high Ba, Sr, K/Rb) with titanite ± rutile partly prevailing in the residue leading to increase in Nb/Ta ratio (Fig. 13) and causing a Nb–Ta minimum (Fig. 10). Archean TTGs can also have high Nb/Ta ratios (e.g., John et al., 2011) and thus such source rocks are also possible contributors for the Nb/Ta variation. The deviation towards negative εNd (T) values of A60 favors a considerable contribution from the Archean crust (Table 1; Fig. 6a). The large variation of zircon εHf (T) values in the AC granitoids (A24, A25, A60) is interpreted by incomplete mixing of dominant Paleoproterozoic (− 2 to − 6) source component with minor Archean (− 14 to − 20) source component (Fig. 6b; Appendix C). Lower crustal xenoliths have been studied from the western kimberlite field (Fig. 1). Hölttä et al. (2000) estimated crystallization at ca. 800– 900 °C and at 22–38 km crustal depths for the xenoliths using mineral core compositions, but Peltonen et al. (2006) argued that these

Fig. 13. Nb/Ta–Zr/Hf relation for samples of this study. Major silicate reservoirs modified from John et al. (2011) and Münker et al. (2003). Dashed lines represent chondritic values for Nb/Ta and Zr/Hf. AC — Archean continent; AP — Archean-Proterozoic boundary; AP/A — A-type granitoids from the AP boundary; CFGC — Central Finland granitoid complex; CFGC/A — A-type granitoids from the CFGC; BZ — boundary zone (BZ); and BB — Bothnian Belt.

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xenoliths are from variable depths from the deep, dense mafic highvelocity layer (35–60 km; Korja et al., 1993). The lower crustal xenoliths are dominated by hydrous mafic garnet granulites with minor felsic granulites and gabbros without any eclogites, but Kukkonen et al. (2008) concluded that a small proportion of crustal eclogites are present in lower part of the high-velocity layer based on kimberlite-hosted garnet xenocrysts. Some crustal xenoliths show a variation from Archean (up to 3.5 Ga) to Paleoproterozoic and some only Paleoproterozoic (ca. 1.9–1.7 Ga) zircon U–Pb ages with main peaks at 1.90–1.85 Ga and 1.81–1.78 Ga (Peltonen et al., 2006). These zircon age peaks coincide well with the ages of the AC magmatism (Huhma, 1986; this study) and the shoshonitic lamprophyres (Woodard and Huhma, 2015; Woodard et al., 2014), respectively. Hydrous mafic granulites have εNd (1.9) values between +3.9 and −2.7 and gabbros show variation from Archean with εNd (1.9) value of −8.7 to Paleoproterozoic εNd (1.9) value of +1.4 and TDM of 2180 Ma (Peltonen et al., 2006). The mafic granulite xenoliths are typically amphibole-bearing and hydrous and they are enriched in Ba and Rb, and have low K/Rb ratios of 150–400, which features are not typical for lower crustal mafic xenoliths (Hölttä et al., 2000). The Sr values are not high and typically the xenoliths show small Eu minimum in their chondrite-normalized REE patterns (ibid). These geochemical characteristics and the on average more juvenile origin of the mafic granulite xenoliths do not favor that they would represent the source for the AC felsic melts. The proposed sources for the AC felsic granitoids are melting of the 2.1– 2.05 Ga plume-derived (Hanski and Huhma, 2005; Lahtinen et al., 2015) magmatic underplate (Fig. 12) and assimilation/melting of the Archean lower crust. The A-type granitoids from the AP (AP/A) show a variable sedimentary component as discussed above. Apart a sample with very high GdN/ YbN ratio (Fig. 14) the AP/A granitoids show low Sr/Y (Figs. 8 and 14) indicating plagioclase in the melt residue and/or very low Sr in the sedimentary source. Based on the GdN/YbN ratios the AP/A granitoids can be divided to “normal” and high GdN/YbN granitoids where the latter also plot towards elevated Nb/Ta ratios in Fig. 13. This could favor a mixed eclogite/amphibolite and sedimentary source for these AP/A granitoids. The AP I-type granitoids are also variable in composition and few of them show a strong sedimentary component (Fig. 9). Quartz diorite A105 and tonalite A1968 are rather similar in composition (Appendix A; Fig. 14) but the Sm–Nd data (Table 1 and Fig. 6) require different protolith ages. High K/Rb, Nb/Ta, Sr/Y and GdN/YbN (Figs. 8, 10, 13 and 14) in many AP granitoids favor a mafic source with eclogitic and/or garnet amphibolitic residue. Both the I-type CFGC and CFGC/A granitoids follow the silicate differentiation trend in the Nb/Ta–Zr/Hf diagram where the one high Nb/Ta sample (Fig. 13) is highly differentiated siliceous granite. The I-type intermediate granitoids have variable but in general higher Sr/Y ratios than the CFGC/A granitoids and typically a trend to lower GdN/YbN ratios with decreasing Sr/Y, favoring feldspar fractionation (Figs. 8 and 14). The mafic and intermediate I-type plutonic rocks are interpreted as formed from mixing of mantle-derived and crustal melts where the felsic granite end-member represents crustal melts deriving from fertile K-rich intermediate crust (Lahtinen, 1994; Lahtinen and Huhma, 1997; Nironen et al., 2000). The input of mafic magma into upper crustal magma chamber is possibly seen in the granodiorite A1973 where zircon core-rim differences in δ18O values show constantly lower values in rims reflecting the mantle-derived input. The TDM variation between 2.11 and 2.42 Ga for the CFGC granitoids (Table 1 and Fig. 6) favors the occurrence of older crust (ca. 2.0–2.1 Ga) as proposed earlier (Lahtinen, 1994; Lahtinen and Huhma, 1997; Rämö et al., 2001). The eastern CFGC/A granitoids are dominated by intermediate granitoids and they belong to type 3b post-kinematic plutons (Nironen et al., 2000), whereas the CFGC/A granitoids in the western part of the CFGC comprise also felsic variants and belong to the type 2 granitoids (ibid.). The major difference of the eastern CFGC/A granitoids to the AP/A granitoids is the higher K/Rb ratios (Fig. 8) and less aluminous

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Fig. 14. Eu*/Eu versus GdN/YbN and Sr/Y versus GdN/YbN diagrams for granitoids from this study. SmN, EuN, GdN, YbN – chondrite normalized values (Boynton, 1984). Eu*/Eu = (EuN) / (SmN)0.5 × (GdN)0.5. See Fig. 13 for legend.

nature (Fig. 9) in the former. Nironen et al. (2000); see also Elliott et al., 1998; Rämö et al., 2001) proposed that the extraction of I-type melts left a compositionally variable hornblende and biotite-poor granulitic residue in the lower crust. Elliott (2003) studied one type 3 and one type 2 post-kinematic plutons from the CFGC (south of the study area in Fig. 2) and suggested that parental magma for these granitoids formed by partial melting of recently emplaced, mantle-derived, mafic rocks mixed with variable proportions of anatectic melts of mafic granulite. The very homogeneous Nd isotopic composition (Fig. 6; Rämö et al., 2001) with typical TDM ages on the order of ~2.2 Ga favor a rather similar protolith age for both the CFGC I-Type and CFGC/A granitoids and excludes a very large contribution from a juvenile source in their genesis. The BZ granitoids have typically higher K/Rb and Sr/Y and younger TDM ages, between 1.94 Ga and 2.16 Ga, than the adjacent CFGC granitoids (Figs. 6, 8 and 14). All these favor a more mafic and, at least, more juvenile crustal source compared to the CFGC I-type granitoids. The 1.92 Ga granodiorite A292 is similar in age and composition to the juvenile gneissic tonalites and granodiorites found in the Older Svecofennia (Fig. 2), which were interpreted to represent low degrees of melting of low-K tholeiitic island arc basalts with gabbroic to amphibolitic residues (Lahtinen, 1994). The low Sr/Y and deep negative Eu anomaly favor large amounts of plagioclase in the source residue and the low GdN/ YbN ratios indicate that garnet or orthopyroxene was not abundant in the source residue (Fig. 14). Samples from younger (1.90 Ga), also juvenile (Table 1; Fig. 6), granodiorite A1270 deviate to higher Al2O3 (Appendix D) and have high Sr/Y and GdN/YbN (Fig. 14), which all fit with plagioclase-poor garnet granulitic/eclogitic residue indicating relatively deep melting level. 5.4. Magma sources, magma chambers and crustal profile along FIRE 3A The challenge in studying the deep crust is the lack of direct evidence in the form of samples apart from kimberlite xenoliths. Granitoid magmas can have a complicated and variable genesis but anyhow they give us indications of their sources deeper in the crust. Seismic reflection method is capable of providing us a snapshot from present crustal architecture down to the depths of Moho (e.g. Clowes, 1993; Ehsan et al., 2014; Goleby et al., 2004; Kashubin and Juhlin, 2010). Abrupt changes of acoustic impedance (product of density and seismic velocity) cause seismic waves to reflect, and this phenomenon is utilized in the seismic surveys that are usually designed to image contacts between different rock types. Typically crustal scale seismic data is acquired along existing roads. Thus seismic profiles are not perpendicular to the geological structures and only apparent dips of the structures are observed. In general, surface seismic reflection method is best suited for subhorizontal features and steep structures are not directly

imaged. Clearly 2D profiles cannot solve complicated 3D structures typical for the deformed and metamorphosed bedrock areas, and certain continuity and contrast in acoustic impedance is required in order to image a certain feature with seismic data. In the following, we try to incorporate the results of our and previous studies along with the FIRE 3A crustal scale seismic reflection profile to locate different magma source regions and magma chambers. In Fig. 15 the Fire 3A profile is shown viewed from the south to north. Selected major tectonic boundaries and faults are interpreted along with the inferred dips of the reflective units. These dip directions are inferred in the 3D modeling environment by combining surface geology and aeromagnetic trends with the true location of the seismic profile and in the different apparent dips of reflectors along the profile. The metamorphism along the profile at present crustal level varies from typical low-P amphibolite facies to local granulite facies metamorphism with pressures at 4–6 kb (Chopin et al., 2014; Korsman et al., 1999). For simplification, we assume that the present exhumation level is a 15 km deep cut from the situation at 1.89–1.85 Ga when most of the studied granitoids intruded. Elliott et al. (1998) interpreted that the late amphibole crystallization, presumably at 10–15 km, in granitoid A218 (Fig. 2) could register conditions at the level of emplacement, which is in line with metamorphic data. The upper crustal magma chambers, now represented by exposed plutons and batholiths (P in Fig. 15) would have their lower floors 5–10 km deeper in the crust, based on the non-reflectivity below the plutons. Most of the kimberlite xenoliths crystallized at ca. 800–900 °C and at 22–38 km crustal depths (Hölttä et al., 2000), but up to 50 km crustal depth has been calculated for one sample (Järvinen, 2015; unpublished M.Sc thesis). As discussed in Section 5.3 the mafic granulite xenoliths are not likely the source for the AC melts, and we propose that the xenoliths derive from magma channels (MC in Fig. 15) of the AC bimodal magmatism. The same channels have been later used by the 1.8 Ga and kimberlite magmas. We suggest that the AC felsic melts (1.87– 1.85 Ga) derive for a large part from partly eclogitic, garnet amphibolites formed from the 2.1–2.05 Ga magmatic underplate. Examples of this magmatism are preserved as the Siilinjärvi mafic volcanism, and as abundant mafic feeder dikes (Vuollo and Huhma, 2005) on the present crustal level. The Archean lower crust contributed via assimilation and incremental melts to create the variable compositions of the felsic AC melts. The AP I-type tonalites and granodiorites are heterogeneous in composition and show different degrees of sediment assimilation. Some samples, like tonalite A1968, indicate derivation from a deep mafic source (Fig. 14) which can be a rather juvenile lower mafic crust or a subducting slab. The AP/A granitoids have a dominant sedimentary component, which is interpreted to be caused by melting of already migmatized sedimentary rocks ± amphibolites in different depths in


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Fig. 15. Migrated seismic section for Fire 3A (Fig. 2; Kukkonen et al., 2006; Sorjonen-Ward, 2006) is shown in 3D looking from the south to north. Selected major tectonic boundaries (solid lines), faults (dashed lines) and inferred dip directions (arrows) of the reflective units are indicated. The dip directions are interpolated combining surface geology and aeromagnetic trends with the seismic reflection profile in 3D. Changes in the apparent dip are due to changes in the profile orientation in relation to the strike of geological structures. In the middle part of the profile also data from FIRE 1 (Fig. 1; Kukkonen et al., 2006) was used for interpretation. Blue arrow — dip direction of the 1.88–1.87 Ga stacking affected by 1.87–1.86 Ga buckling; Green arrow — dip direction of 1.83–1.81 Ga stacking. Red solid lines in the AP-AC are inferred early (ca. 1.91 Ga) thrusts. P — plutons/batholiths; MC — magma chambers; m — migmatization; Bold capital case letters indicate the source areas for crustal melts discussed in text. C/I — CFGC I-type; C/A — CFGC A-type.

the middle crust. The sedimentary input may have occurred in several stages as seen in higher δ18O values in the zircon rims than in the cores in some plutons. The AP boundary granitoids have zircon U–Pb ages between 1887 Ma and 1879 Ma and a median age 1882 Ma, which is interpreted as the main stage of both types of magmatism. Younger ca. 1875 Ma ages from mafic plutonic rocks are found north of the FIRE 3A profile (Fig. 2). Traditionally, the CFGC granitoids have been divided to foliated synkinematic (1.89–1.88 Ga) and less deformed post-kinematic (1.88– 1.87 Ga) granitoids (Nironen et al., 2000). The CFGC I-type granitoids of this study vary typically in age between 1891 Ma and 1883 Ma, and the CFGC/A granitoid A85 has an age of 1883 ± 17 Ma with a large error and CFGC/A granitoid A1972 has an age of 1876 ± 6 Ma (Table 1). There seems to be a slight tendency that the I-type granitoids are older than 1885 Ma and that the A-type granitoids are typically ca. 1880 Ma (see Fig. 2 and Rämö et al., 2001) but both A447b (I-type) and A1972 (CFGC/A) have a same age at 1875 Ma. It is suggested that the I-type granitoids can be divided into two groups at ≥1885 Ma and ≤1882 Ma, where the latter ones overlap in age with the CFGC/A granitoids. In Fig. 15, we suggest that the CFGC I-type granitoids often derive from slightly deeper source than the CFGC/A granitoids and that, at least partly, the CFGC/A granitoids derive from melting of an I-type intermediate source. The BZ granitoids are similar in age to the CFGC I-type granitoids (Table 1) but have in general a deeper origin (Figs. 14 and 15) and derive from a more juvenile source (Fig. 6). The granodiorite A292 (1921 ± 5 Ma) from the BB is a pre-thickening granitoid, correlated with similar granitoids in the older Svecofennia (Lahtinen, 1994; Lahtinen and Huhma, 1997) and derive from melting of a juvenile mafic source. Granodiorite A1270b show similar juvenile mafic source, but is younger (1898 ± 2 Ma), and the melting has now occurred in depth where plagioclase was a minor component or absent. Both slab melting or melting of a thickened crust are possible. 5.5. Tectonic implications Models that have been proposed to account for the present crustal image of the central and western part of the FIRE 3A all include crustal scale stacking/thrusting, but the importance and degree of extension vary from minor (Sorjonen-Ward, 2006), moderate (Kukkonen et al.,

2008) to large-scale gravitational spreading (Korja et al., 2009; Nikkilä et al., 2015). Based on granodiorite A1270b (see above), the stacking could have started already 1.90 Ga ago but the common strong foliation seen in the 1890–1885 Ma granitoids (Nironen, 2003; Nironen et al., 2000) favors that stacking/thrusting has continued at least until 1885 Ma. Sorjonen-Ward (2006) proposed two models where either accretion of W-facing magmatic arc at 1.89–1.87 Ga or later postmagmatic crustal scale imbrication with NNW vergence at 1.84–1.80 Ga would have caused the presently seen crustal-scale reflective bands (Fig. 15). We present here a tentative model for testing. The first stage is the collision at c. 1.91 Ga between the Archean continent and the Paleoproterozoic microcontinent-arc collage (e.g. Gaál, 1990; Kohonen, 1995; Lahtinen et al., 2005, 2009, 2015; Nironen, 1997). A shallow slab breakoff occurred followed subsequently by ophiolite obduction and a formation of a foreland fold-and-thrust belt. Mantle– crust decoupling of the Archean lithosphere occurred (Peltonen and Brügmann, 2006) and a crocodile jaw formed, where the arc rocks are partly buried beneath the bounding Archean continent (Fig. 12; Lahtinen et al., 2015 and references therein). The collision was followed by a subduction polarity reversal and subduction under Keitele ribbon (CFGC) starting at 1.90 Ga followed shortly by a collision that initiated at ca. 1.88 Ga, forming the Fennian orogen (Lahtinen et al., 2005, 2014). Both collisions had about similar SWW-ENE shortening directions and caused the crustal-scale stacking of both extended continent margin and the Keitele continental ribbon. We suggest that the present lower crust under the CFGC is composed of ca. 2.0 Ga melt-extracted granulites and mafic cumulates (see Annen et al., 2006, 2015) formed during 1.90–1.88 Ga arc magmatism. The crustal thickness could have been in the order ca. 20–40 km (A292 and Older Svecofennia; Kukkonen et al., 2008) before the first stacking stage at 1.91 Ga and then around 40–50 km during the 1.90–1.88 Ga arc stage followed by the stacking stage at ca. 1.88 Ga and then 65–75 km (cf. Nikkilä et al., 2015) before the ≤ 1.80 Ga exhumation. The ≤1.88 Ga AP granitoids are considered to represent the end of compression/transpression and change to neutral or transtensional stage (see also Nironen, 2003). Lahtinen et al. (2014) proposed that 1.87– 1.86 Ga buckling caused the formation of the coupled Bothnian oroclines (Fig. 1). The buckling mechanism (Johnston et al., 2013) implies a space problem that was solved in the coupled Bothnian oroclines

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by proposing material flow in middle/lower crust and mantle (ibid.). One possibly is that the ≤ 1875 Ma granitoids in the CFGC are syn-buckling granitoids, intruding to open spaces created by the onset of buckling. The shear zones in the AP boundary (Fig. 15) have been active, at least from 1.87 Ga onwards and the magmatic age peak at 1870– 1865 Ma in the AC includes large plutons that have intruded accompanied by strike-slip deformation (Gaál and Rauhamäki, 1971; Halden, 1982; Sorjonen-Ward, 2006). This would indicate at least 10–20 Ma active horizontal movement during buckling. The deep nature of the fault zones in the AP (Fig. 15) is seen in the abundant occurrence of primitive mantle-derived magmas and related Ni-deposits (Makkonen and Huhma, 2007), reflecting rapid ascent along faults to middle-upper crust. Major orogenic overprinting occurred at 1.83–1.81 Ga during NW–SE convergence (cf. Sorjonen-Ward, 2006) contracting the southernmost limb of the Bothnian Orocline and partly overthrusting it by the rocks of Southern Svecofennia (Appendix D; Lahtinen et al., 2014). The interpreted upper crustal NW vergent stacking (green units in Fig. 15) of the CFGC is related to this stage. The above presented model elegantly explains the preservation of thick crust and mantle (Fig. 12) to buckling and orocline formation (see Johnston et al., 2013) later modified by younger contraction. In this model most of the heat needed to produce crustal melting is related to mantle-derived magmas and thus this model differs from the conductive heat model for the CFGC by Kukkonen et al. (2008). This model lacks large scale extension and thus differs from the core complex model by gravitational spreading (Korja et al., 2009; Nikkilä et al., 2015). Our model is based on the buckling mechanism for generating the Bothnian Oroclines and the equidimensional CFGC, and future studies will shed light on how the Bothnian Oroclines have formed. 6. Conclusions We present a new interpretation of the seismic reflection profile FIRE 3A imaging the present crustal structure from the Karelia province (Archean + Paleoproterozoic cover) to Paleoproterozoic Svecofennia province in Fennoscandia. The profile has been further divided into Archean continent (AC), Archean-Proterozoic boundary (AP), Central Finland Granitoid Complex (CFGC), boundary zone (BZ) and the Bothnian Belt (BB). The 1.92–1.85 Ga plutonic rocks along the profile have been studied using geochronology, geochemistry and isotopic data (Sm–Nd, Lu–Hf, O) to assess the composition of crust and subcontinental lithospheric mantle (SCLM). Granitoids from the AC have typically long, euhedral and zoned zircons, whereas the zircons from other granitoids are typically more equidimensional. Apatite occurs as a main phase of inclusion in many zircons but quartz, orthoclase, amphibole and biotite inclusions are also found. Main exception is granite A545 where the main phases of zircon inclusions are compound minerals and undefined feldspar-Femineral-mix, interpreted as hematite micro-inclusions within the feldspar, often surrounded by monazites and xenotime as crusts around inclusion nodules. Measured core-rim pairs in zircons indicate in some cases clear differences in δ18O values interpreted as due to new mantle-derived magma pulse entering the magma chamber (lower δ18O rim values) or due to sedimentary assimilation (higher δ18O rim values). The AC plutonic rocks (1.87–1.85 Ga) form a bimodal suite where the proposed mantle source for the AC mafic rocks is the 2.1–2.0 Ga metasomatized lower part of the Archean SCLM. Further contamination by continental crust before entering the middle–upper crust is considered responsible for the intermediate composition, “subduction-like” Nb–Ta minimum in spidergrams, and negative εNd (T) values. The felsic AC granitoids show variable source components, e.g. seen in Sm–Nd and Lu–Hf data, but for most granitoids a low degree melting of eclogitic or garnet-bearing amphibolites from an enriched mantle source (high Ba, Sr, K/Rb) with titanite ± rutile partly prevailing in the residue (Nb/

Ta–Zr/Hf relation) is proposed. The suggested source is the 2.1– 2.05 Ga plume-derived magmatic underplate and the differences in the εNd and zircon εHf (T) values are due to variable degree of assimilation/melting of the Archean lower crust. The AP comprises of mafic and ultramafic rocks of which some plot in the enriched mantle array in the Th/Yb–Ta/Yb diagram. Typically the AP plutonic rocks (ca. 1.88 Ga) seem to be characterized by a variable sedimentary component seen, e.g., in the AP A-type granitoids (AP/A) which follow the sediment assimilation trend in ASI diagram, have high δ18O values in zircons and exhibit negative Ba anomalies (Rb–Ba–Th in spidergram), similar to those in sedimentary rocks. A mixing/assimilation of enriched mantle-derived melts with melts from already migmatized sedimentary rocks ± amphibolites are proposed for the origin of AP/A granitoids. Some AP I-type tonalites indicate derivation from a deep mafic source which can be rather juvenile lower mafic crust or a subducting slab. The CFGC is characterized by abundant I-type and A-type (CFGC/A) intermediate and felsic granitoids and scarcity of mafic rocks with SiO2 ≤ 52%. The I-type granitoids are divided into two groups at ≥ 1885 Ma and ≤ 1882 Ma, where the latter ones overlap in age with the CFGC/A granitoids. Both I-type CFGC and CFGC/A granitoids are interpreted to have formed from mixing of Paloeproterozoic SCLMderived melts with crustal melts from hydrous and dry intermediatefelsic igneous sources, respectively (see also Elliott et al., 1998; Lahtinen and Huhma, 1997; Nironen et al., 2000). The igneous-affinity geochemistry, dominantly magmatic δ18O values in zircons and TDM variation between 2.11 and 2.42 Ga of the CFGC granitoids favor the occurrence of older igneous crust (ca. 2.0 Ga) (see Lahtinen, 1994; Lahtinen and Huhma, 1997; Rämö et al., 2001) in the genesis of the CFGC granitoids and excludes the marginal basin hypothesis requiring melting of large amounts of sedimentary rock (Rutland et al., 2001, 2004; Williams et al., 2008). The BZ granitoids are similar in age to the CFGC I-type granitoids but have typically higher K/Rb and younger TDM ages, between 1.94 Ga and 2.16 Ga, than the adjacent CFGC granitoids which favor a more juvenile crustal source for the BZ granitoids. The 1.92 Ga granodiorite in the BB is similar in age and composition as the juvenile gneissic tonalites and granodiorites from the Older Svecofennia in the AP boundary, and a similar shallow melting (plagioclase stable) of juvenile mafic source is proposed. Slightly younger granodiorite (1.90 Ga) from the BB is also from juvenile source but now having plagioclase-poor garnet granulitic/eclogitic residue indicating relatively deep melting level. A tentative tectonic model based on Lahtinen et al. (2014, 2015) is discussed. We suggest that the present high-velocity lower crust under the CFGC is composed of melt-extracted igneous granulites (source age ca. 2.0 Ga) and mafic cumulates which both formed during 1.90–1.88 Ga arc magmatism. The crustal thickness has been in the order ca. 20–40 km before the first stacking stage at 1.91 Ga and then around 40–50 km before the new stacking stage at ca. 1.88 Ga. The ≤1.88 Ga ages in the AP and CFGC granitoids is considered to represent the end of compression/transpression and change to neutral or transtensional stage before the 1.87–1.86 Ga buckling that caused the formation of the coupled Bothnian Oroclines. The shear movement along the AP boundary is related to buckling and the magmatic age peak at 1870–1865 Ma in the AC includes large plutons that have intruded accompanied by this strike-slip deformation. Major orogenic overprinting occurred at 1.83–1.81 Ga during NW–SE convergence contracting the southern limb of the southern Bothnian Orocline and partly overthrusting it by the rocks of Southern Svecofennia. This stage is seen in the upper crustal NW vergent stacking in the CFGC. This tentative tectonic model relates the preservation of thick crust and mantle to buckling and orocline formation, later modified by younger contraction. This model lacks large scale extension and thus differs from the core complex model by gravitational spreading (Korja et al., 2009; Nikkilä et al., 2015) and the heat production is related to


magmatic additions and not to conductive warming as proposed by Kukkonen et al. (2008). Future studies will hopefully solve this paradigm. Supplementary data to this article can be found online at http://dx. doi.org/10.1016/j.lithos.2016.07.014. Acknowledgements We thank Tuula Hokkanen, Leena Järvinen, Arto Pulkkinen and Hugh O'Brien for their help in the laboratory. The Nordsim ion-microprobe facility was supported by the Natural Science Funding Agencies of Denmark, Finland, Norway and Sweden. Lev Ilinsky and Kerstin Lindén are thanked for their invaluable technical assistance. Bernard Bonin and Chris Harris are gratefully acknowledged for their constructive reviews. This is NORDSIM publication number 465.

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