Near-orthogonal deformation successions in the poly-deformed Paleoproterozoic Martimo belt: Implications for the tectonic evolution of Northern Fennoscandia

Accepted Manuscript Title: Near-orthogonal deformation successions in the poly-deformed Paleoproterozoic Martimo belt: Implications for the tectonic evolution of Northern Fennoscandia Author: Raimo Lahtinen Mohammad Sayab Fredrik Karell PII: DOI: Reference:

S0301-9268(15)00289-2 http://dx.doi.org/doi:10.1016/j.precamres.2015.09.003 PRECAM 4345

To appear in:

Precambrian Research

Received date: Revised date: Accepted date:

26-5-2015 17-8-2015 4-9-2015

Please cite this article as: Lahtinen, R., Sayab, M., Karell, F.,Near-orthogonal deformation successions in the poly-deformed Paleoproterozoic Martimo belt: Implications for the tectonic evolution of Northern Fennoscandia, Precambrian Research (2015), http://dx.doi.org/10.1016/j.precamres.2015.09.003 This is a PDF file of an unedited manuscript that has been accepted for publication. As a service to our customers we are providing this early version of the manuscript. The manuscript will undergo copyediting, typesetting, and review of the resulting proof before it is published in its final form. Please note that during the production process errors may be discovered which could affect the content, and all legal disclaimers that apply to the journal pertain.

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Near-orthogonal deformation successions in the poly-deformed Paleoproterozoic Martimo belt: Implications for the tectonic evolution of Northern Fennoscandia Raimo Lahtinen*, Mohammad Sayab, Fredrik Karell Geological Survey of Finland, P.O. Box 96, FI-02151 Espoo, Finland *Corresponding author e-mail: [email protected]

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16 Abstract:

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Structural mapping integrated with aeromagnetic and Anisotropy of Magnetic

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Susceptibility (AMS) data form complimentary and robust tools for analysis of regional

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structures and fabric overprinting relationships. This multi-scale approach is used to

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constrain the geometry and tectonic evolution of sparsely exposed and flat-lying

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outcrops of the Martimo belt (MB) of northern Fennoscandia. The MB records a history

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of near-orthogonal deformations upper greenschist to lower amphibolite facies

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metamorphism between ca 1.92-1.79 Ga. The earliest deformation (D1) involved east-

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directed thin-skin thrusting, and the development of recumbent F1 folds and a N-S

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trending S1 fabric. D2 deformation is prominent in the central part of the MB and is

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characterized by N-S shortening with the development of E-W trending F2 folds, an

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associated pervasive steeply-dipping S2 foliation and type-2 fold interference pattern.

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The MB and Mellajoki suite exhibit regional scale D2 synformal and antiformal

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refolded structures, respectively. D3 is more heterogeneously developed and manifests

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bimodal orientations with NNW-SSE and WNW-ESE structural trends. It is pronounced

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in the central and eastern parts of the MB but weakly developed in the western part. The

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steep dip of S3 in the region suggests that it formed in response to NE-SW bulk

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shortening. D4 structures have a conspicuous NNE-SSW trend that is visible as

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lineaments on the aeromagnetic image and locally in the field. The last deformation

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event (D5) affected almost the entire MB and was characterized by broadly WSW-ENE

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oriented crustal shortening. The N-S oval-shape geometry of the eastern part of the MB

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is attributed to the D5 event. Contrasting with well developed overprinting relationships

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between D1 to D3 structures, the relative timing of D3-D5 could not be established in

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the study area, but can be inferred from structures with similar trends in the Central

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Lapland Granitoid Complex, which is located north of the MB. The D1 to D5

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deformation scheme can be readily reconciled with the tectonic evolution of the

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Lapland-Savo, Fennian, Svecobaltic and Nordic orogens of the composite Svecofennian

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orogen. The age of D1 in the northern part of the Lapland-Savo orogen is ≤1.91 Ga

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based on crosscutting granitoids and maximum deposition ages of sedimentary rocks.

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D2 (1.90-1.89 Ga) has been linked to orogen parallel shortening during the late stage of

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southern part of the Lapland-Savo orogen. Several orthogonal tectonic events in

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northern Fennoscandia indicate amalgamation of crustal blocks around the core of the

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Archean Karelia continent at 1.92-1.77 Ga, and that Fennoscandia was a key component

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of the Columbia/Hudsonia/Nuna supercontinent.

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Research Highlights

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- The Martimo Belt (Fennoscandia) manifests complex Paleoproterozoic tectonic

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evolution

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- Five near-orthogonal deformation successions (D1 to D5) are identified

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- Each tectonic event is characterized by distinct bulk crustal shortening

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- Strong strain heterogeneity is partly due to the Archean basement involvement

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Keywords: Svecofennian orogen, Fennoscandia, Paleoproterozoic, Folds, AMS,

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Aeromagnetic data

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1. Introduction

63 One of the major difficulties in the structural analyses of polydeformed Precambrian

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terranes, especially cratonized flat-shield areas, is the scarcity of vertical exposures of

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the bedrock. This well known problem is in the Fennoscandian shield region further

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exacerbated by abundant Quaternary sedimentary cover. Generally, defining the

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regional-scale tectonic architecture of such ancient orogens is subject to considerable

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uncertainty, due to limited exposures and complex polyphase structures both at local

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and regional scale (e.g. Sayab, 2009). Despite the poor outcrop, the Fennoscandian

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shield represents a window into a root zone of a Paleoproterozoic orogen, which makes

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it an ideal natural laboratory to investigate the interplay between deformation and

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metamorphism in ancient times of Earth's history when the processes of lithospheric

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deformation and metallogeny were not necessarily identical to the ones presently active

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(e.g. Riller et al., 2010; Cagnard et al., 2011). Keeping in view the mapping limitations

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in such terrains, advanced structural geology, petrofabric and geophysical methods

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provide a robust and systematic workflow to examine key areas at scales that can be

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extrapolated to large regions. In this study, we have integrated quantitative structural

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field data with the Anisotropy of Magnetic Susceptibility (AMS) measurements on

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oriented drill cores, and high-resolution (survey flights at 30 m altitude, 200 m line

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spacing) regional aeromagnetic images to model the tectonic evolution of the multiply

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deformed Martimo metasedimentary belt (MB) of Northern Finland. The combination

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of three independent methods significantly helped us in developing the tectonic

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evolution of MB from scratch, and especially the AMS data was highly useful in

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recognizing earlier ductile structural imprints. This paper illustrates the usefulness of

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combining aeromagnetic, AMS and structural field data to constrain the geological

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evolution of Precambrian shields based on analysis of limited outcrop of

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metasedimentary sequences.

89 The MB constitutes a part of the Peräpohja Belt, which lies in the centre of the

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Fennoscandian shield, and thus holds a key position to reconstruct its complex

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Paleoproterozoic tectonic evolution of the region (Fig. 1). It strikes roughly east-west in

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typically flat land with Quaternary sedimentary cover only leaving sporadic exposures

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of mainly steeply-dipping Paleoproterozoic rocks. The belt hosts relatively incompetent

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rock units, such as metapelites interbedded with metapsammites (Kaskimaa and

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Ristivuoma metasedimentary rocks, Fig. 2a), and thus preserves a sequence of tectonic

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events that are not fully developed elsewhere in the region. Metamorphic grade in the

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MB increases from west to east, and is generally characterized by upper greenschist

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(biotite) to lower amphibolite facies (biotite+cordierite±andalusite), respectively. Most

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of the outcrops have been mapped in the 1970's and 1980's as a part of the Finnish

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Geological Survey national bedrock mapping programme at 1: 100,000 scale, which

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provides the basis for our more detailed structural analysis. We have recently revisited

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some of these locations and selected 32 outcrops (Fig. 2a), where heterogeneously

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developed structural sequences show systematic changes variation in their orientation.

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Our structural survey indicates that strain is kinematically partitioned into map-scale

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domains, such that individual outcrops typically preserve only two and a maximum of

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three deformation phases of a total of five that affected the study area. We show that

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this complex structural evolution of the MB fits nicely with the overall tectonic

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evolution of northern Fennoscandia, and allows us to refine the still poorly constrained

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Paleo-Mesoproterozoic supercontinent reconstructions (e.g. Meert, 2012). Furthermore,

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understanding the tectonics of the MB as part of the Peräpohja Belt, is important from

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an economic perspective as the area hosts potential Cu-Au occurrences of various

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settings (Eilu, 2012) and has recently received attention from mining companies.

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2. Geological setting

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The main Paleoproterozoic orogenic evolution of Fennoscandia (Fig. 1 inset) has been

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divided into the Lapland-Kola orogeny (1.94–1.86 Ga; Daly et al, 2006) and the

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composite Svecofennian orogeny (1.92–1.79 Ga; Lahtinen et al, 2005; 2009). The latter

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is further divided into the Lapland-Savo, Fennian, Svecobaltic and Nordic orogenies.

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Whereas the Lapland-Kola orogen shows only limited formation of new crust, the

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composite Svecofennian orogen forms a large volume of the Paleoproterozoic crust in

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the central and southwestern part of the Svecofennian province. The equidimensional

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Fennia orogen in the central part of the Svecofennian Orogen has recently been modeled

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to have formed by the buckling of a linear orogen about vertical axes of rotation into

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coupled Bothnian oroclines (Lahtinen et al., 2014). The evolution of the southern and

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southwestern parts of the Svecofennian Orogen have been also interpreted with a

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tectonic switching/accretionary orogen model (Hermansson et al., 2008; Saalman et al.,

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2009; Bogdanova et al., 2015). Bogdanova et al. (2015) further concluded that the

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accretion was interrupted at 1.82–1.80 Ga by the oblique continent-continent collision

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of Fennoscandia and Volgo-Sarmatia (the Svecobaltic orogeny in this study).

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The northern part of Fennoscandia is characterized by an approximately N−S trending

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Paleoproterozoic Pajala shear zone (Kärki et al., 1993), also named as the Baltic-

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Bothnian megashear (Berthelsen and Marker, 1986), comprising several subvertical

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fault sets (Fig. 1). This zone has been earlier interpreted as an intracontinental strike-slip

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shear zone (Berthelsen and Marker, 1986), an intracontinental rift that subsequently

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inversed to a thrust and fold belt (Nironen, 1997) or a suture (a plate boundary or

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transform fault) (Lahtinen et al., 2005; 2009). Lately Lahtinen et al. (2015a) revised the

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latter model based on new geochronological data. Their model proposes that the two

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Archean continental units (Norrbotten and Karelia) collided at ca. 1.92-1.91 Ga, where

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the Kittilä allochthon (arc part) was obducted upon the Karelia continent while a

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foreland fold and thrust belt formed. Thus, a cryptic suture between the Norrbotten and

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Karelia plates within the Lapland-Savo orogen is inferred (Fig. 1). This suture was

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multiply reactivated after continental collision with both lateral and vertical movements,

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and thus the Pajala shear zone proper is mainly a younger overprinting feature. The

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occurrence of a sharp plate boundary is also favored between two Archean mantle

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lithospheres (Norrbotten and Karelia), reflecting different fossil anisotropic structures

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(Plomerova et al., 2011; Vecsey et al., 2014).

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In the Norrbotten province, west of the cryptic suture (Fig. 1), the bedrock is dominated

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by medium to high-grade metamorphosed sedimentary rocks, mainly pelitic and arkosic

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to quartzitic types, with subordinate intermediate to mafic metavolcanic rocks. Close to

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the cryptic suture, the supracrustal rocks are often strongly recrystallized and,

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particularly in the case of the metapelites, migmatitic in nature (e.g. Bergman et al.,

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2001; Jonsson and Kero, 2013). In Fig. 1 these supracrustal rocks have been divided

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into two classes, one dominated by ≤1.91 Ga rocks and other by undefined, mainly

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Paleoproterozoic, rocks (see Lahtinen et al., 2015a). Plutonic rocks have not been

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differentiated but include felsic to mafic intrusives representing three age groups: ca.

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1.89-1.87 Ga, ca. 1.88-1.87 Ga and ca. 1.81-1.78 Ga (Bergman et al., 2001). The

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bedrock of the Karelia province in Fig. 1 may be divided into four major lithologic

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domains. The southernmost part belongs to the Archean Pudasjärvi Complex, intruded

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by 2.44 Ga layered intrusions. The Paleoproterozoic supracrustal rocks of the Peräpohja

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Belt were deposited in both continental and marine environments on top of the

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Pudasjärvi Complex. The supracrustal rocks are normally well preserved and have been

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metamorphosed in greenschist to lower amphibolites facies conditions. The Central

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Lapland Granitoid Complex (CLGC) is poorly studied and consists of upper

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amphibolite facies gneisses and migmatites, and voluminous granitoids, with zircon

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ages ranging from 2.3 Ga to 1.76 Ga (Huhma, 1986; Rastas et al. 2001; Ahtonen et al.

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2007; Lauri et al. 2012). Few inherited zircons 139 and negative epsilon-Nd values of

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granitoids (Huhma, 1986) imply that they derive from the Archean lithosphere. The

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lithology and stratigraphy of the Central Lapland Greenstone Belt (CLGB) have

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similarities with the Peräpohja Belt favoring a common history (see Lehtonen et al.

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1998), probably until the proposed continental breakup at 2.1-2.05 Ga along the present

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western margin of the Karelia province. The break-up was associated with the formation

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of an aulacogen, a failed arm of a triple junction, where the Central Lapland Granitoid

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Complex is considered as a later exhumed root zone of an aulacogen (Lahtinen et al.,

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2015a).

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The structural pattern and evolution of the Northern Fennoscandia is complex and

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largely poorly understood. There are several regional studies on the structural evolution

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of the CLGB (Fig. 1), which either link the evolution to a single progressive event

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(Ward et al., 1989) or to multiple deformation phases (Lehtonen et al., 1998; Evins and

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Laajoki, 2002; Hölttä et al., 2007). Important marker units in the CLGB are the 1.88-

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1.86 Ga supracrustal rocks (Fig. 1), which lack some of the earlier deformation events

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and show only E-W and N-S axial traces (Lehtonen et al. 1998).

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2.2. Geology of the Peräpohja belt and CLGC

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191 The Peräpohja Belt has been studied in detail and two major lithostratigraphic units, the

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Kivalo and Paakkola Groups have been previously established (e.g., Perttunen, 1985;

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Perttunen and Hanski, 2003; Kyläkoski et al., 2012). The ca. 3- to 4-km-thick Kivalo

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Group includes all the 2.5-2.06 Ga supracrustal rocks in Fig. 1 except the Mellajoki

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suite (see Fig. 2a). The Kivalo Group comprises abundant orthoquartzites and arenites,

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conglomerates, basalts, mafic sills and dykes. The traditional definition of the Paakkola

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Group includes the lowermost metasedimentary Martimo Formation, metavolcanic

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Väystäjä Formation (2.05 Ga; Perttunen and Vaasjoki, 2001), bimodal Korkiavaara

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Formation and metasedimentary Pöyliövaara Formation (Hanski et al., 2005; Kyläkoski

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et al., 2012). Lately, all these units have been considered as lithodemic in nature

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(Bedrock of Finland−DigiKP), and are followed as such in this study.

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The Peräpohja Belt is metamorphosed from upper greenschist in the southwest to lower

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amphibolite facies in north to northeastern sectors, (Perttunen and Hanski, 2003). The

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MB is a poly-deformed fold and thrust belt that despite the lack of clear

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lithostratigraphic markers can be differentiated into three lithological types. Graphite

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and Fe sulfide-bearing Liekopalo paraschist occurs in close association with the

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quartzites of the Kivalo Group in the south and east and with the quartzites of the

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Mellajoki suite in the north (Fig. 2a). The second type is corresponds to the Kaskimaa

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greywackes, a thick sequence of turbidites with interlayered pelites characterized by a

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bimodal source of mafic volcanic rocks and quartzites (Lahtinen et al., 2015a). The

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Ristivuoma greywackes are characterized by a more felsic sediment source mainly from

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Paleoproterozoic granitoids as reflected in abundant 1.91-2.0 Ga detrital zircons (Ranta

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et al., 2015; Lahtinen unpublished data). The Väystäjä mafic volcanic rocks are partly

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pillowed with basaltic composition and the associated felsic porphyries have yielded a

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U-Pb zircon age of 2050 ± 8 Ma (Perttunen and Vaasjoki, 2001). A similar maximum

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depositional age has been obtained from a sedimentary unit within an intraformational

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conglomerate inside the Väystäjä mafic volcanics (Lahtinen et al., 2015). The Hosiojoki

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quartz-feldspar gneiss (zircons ca. 1.99 Ga; Lahtinen et al., 2015), north of Väystäjä

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rocks, is correlated with the Korkiavaara arkosite of the Rovaniemi supersuite (Hanski

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et al., 2005). The Mellajoki suite comprises quartzites, mica schists and partly mylonitic

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quartz feldspar gneisses (Perttunen and Hanski, 2003). It is intruded by ca. 2.0 Ga

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Kierovaara granites (Ranta et al., 2015; Lahtinen, unpublished data), which vary from

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foliated granite to augen gneiss where the K-feldspars are strongly stretched. According

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to Perttunen (1991), the Peräpohja Belt, including also the Martimo belt, is a large

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synclinorium consisting of alternating synclines and anticlines with steep E-W striking

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axial planes, indicative of N-S shortening. Unpublished M.Sc theses from small areas

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are available but no regional scale study exists from the Peräpohja belt. Kyläkoski et al.

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(2012) summarized the existing studies in a general sequence of 3 deformations. A first

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deformation phase (D1) of N-S compression would have produced recumbent isoclinal

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folds and a penetrative layer-parallel schistosity. Continued N-S compression produced

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tight upright folds with E-W trending axial planes and crenulation cleavage (D2)

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associated with granitoids aged 1885-1880 Ma (Perttunen and Vaasjoki, 2001; Fig. 1).

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D3 formed by E-W compression and produced open folds with N-S trending axial

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planes but often lacking axial planar cleavage.

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3. Fabric overprinting relationship and tectonic sequences

239 Mesoscopic structural data were collected from selected outcrop localities (Fig. 2a),

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whereas sets of 3 to 4 oriented cores were extracted from 4 locations for AMS analysis

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using a portable hand-held drilling machine. The cores were later re-oriented in the lab

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and cut into 2.5 cm diameter and 2.1 cm long cylindrical pieces. Each core is composed

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of metapelite with distinct biotite lineation. All the measurements were carried out using

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a KLY-3S magnetic susceptibility meter (Agico, Inc.) at the Geological Survey of

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Finland, Espoo (GTK). The technique has widely been used to determine mineral

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lineations in multiply deformed rocks (e.g. Riller et al., 1996; Borradaile and Jackson,

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2010). The purpose of AMS analysis in this study was to determine orientation of

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mineral lineations associated with the D1 event, which are rather poorly exposed or

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restricted due to the limited dimension of outcrops. At all 4 locations, drill core samples

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were extracted from low-strain zones of D2 or D3. Aeromagnetic data for the MB

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integrated with the surface geological map clearly depicts bedding-(sub)parallel

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structures and these were used as a basis for finer scale structural analysis (Fig. 2b). On

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the other hand, detailed mapping of selected locations reveal a markedly heterogeneous

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pattern of deformation characterized by significant variation in the style and intensity of

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folding between different outcrops. The study area was divided into 7 subareas for

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kinematic analysis (Fig. 2a). Below, we illustrate and describe 9 outcrop locations, each

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show two to three distinct fabric overprinting relationships, whereas data and

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interpretations for 16 locations are provided in the electronic supplement (Apx1-Apx16

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and Table A in the Appendices).

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3.1. D1 structures

263 The first deformation episode produced a well developed bedding-parallel cleavage

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(S0/S1), which is particularly pronounced in subarea A, where the metamorphic grade is

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of upper greenschist facies. In this subarea, S1 is steeply dipping, roughly E-W striking

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and axial plane fabric of F1 folds. At location 3 of subarea A, F1 folds are open to

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closed, locally symmetric, (Fig. 3; Apx1) with the fold axes (L01) steeply plunging

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towards the west (lineation and foliation terminology after Bell and Duncan, 1978).

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However, the AMS mineral lineation (AMSL) from the same subarea plunges gently to

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moderately west.

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In subarea B, S1 is roughly N-S striking and crenulated by S2, whereas the AMSL is

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shallowly plunging towards the east (Fig. 4; Apx3-Apx5). In subareas C, D and E (Figs.

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5 and 7; Apx6-Apx12) S1 is generally ENE-WSW and NE-SW striking, whereas in

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subareas F and G (Figs. 8 and 9: Apx13-Apx15), it follows the regional north-south

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trending oval-shape geometry, with local variations. F1 folds are tight to isoclinal at

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locations 16 (Fig. 9) and 17 of subarea F (Apx15) and refolded by D3. The AMSL

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mineral lineation in subarea E is near-vertical (Fig. 7), whereas in subarea F it is

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shallowly plunging either towards the NNW or SSE (Fig. 9).

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3.2. D2 structures

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The S2 foliation is characterized by a pervasive, generally east-west striking and steeply

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to vertically-dipping foliation, and is prominent in and around subareas B, C, and D

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(Figs. 4 and 5; Apx3, Apx5-Apx7). It is associated with F2 folds (Fig. 5) that have open

Page 11 of 49

to closed interlimb angles (~30°-110°). At macroscopic scale and on aeromagnetic

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images, F2 folds are mostly tight, rarely isoclinal, and well outlined by thin black-shale

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horizons in the Martimo belt (Fig. 2a,b). The latter acted as glide planes enhancing

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thrusting towards the east. S2 clearly overprints S0/S1, or is subparallel to it in high

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strain zones of D2 (Figs. 4 and 6; Apx6). At location 23 of subarea C, S2 is near-

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vertical, east-west striking and associated with a shallowly WNW-plunging intersection

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lineation (L0/12) (Fig 5). In the eastern (subareas E, F, and G) and western (subarea A)

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parts of the MB, S2 is faintly developed. The L0/12 intersection lineation plunges

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moderately- to gently towards the west-southwest (see Apx2-Apx5) and west (Apx7 and

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Apx9) in subareas B and C, respectively, but east-northeast in subarea D (Apx10).

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297 3.3. D3 Structures

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D3 is more heterogeneously developed than D1 and D2 structures. The S3 foliation is

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generally NNW-SSE striking, however, WNW-ESE trends are not uncommon. It is

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pronounced in the central and eastern parts of the MB and weakly developed in the

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western part (Fig. 2b). At location 23 of subarea C, S3 is a NNW-SSE striking steeply-

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dipping spaced foliation, whereas the L2-3 intersection lineation is moderately to

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steeply plunging towards the NNW (Fig 5). F3 folds, and associated S3 cleavage, are

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superimposed on F2 folds locally producing type-1 fold interference patterns. F3 folds

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are generally open to closed with interlimb angle typically between 120° and 040°. At

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location 28 of subarea E a NNW-SSE striking S3 foliation, which crosscuts S0/S1, is

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defined by well-aligned < 2 cm elongated cordierite porphyroblasts and associated L3

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mineral lineation (Fig. 7). It is noticeable that bedding-parallel (S0/S1) cordierites are

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coarser-grained than the ones developed along the S3 foliation. In subareas F and G, S3

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overprints S0/S1, whereas F3 folds in both subareas are open with intersection lineation

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(L13) moderately- to steeply-plunging either towards SE or NW, respectively (Figs. 9

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and 10; Ap14-Ap16). Open F3 folds at location 16 (subarea F) buckled F1 folds.

315 3.4. D4 Structures

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D4 structures are generally NNE-SSW trending in the whole Peräpohja region

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(discussed below) and associated with an S4 crenulations cleavage. Despite their

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pronounced geophysical signature as lineaments on the aeromagnetic image (Fig. 2b),

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field evidence of D4 deformation is observed in only a few places. At location 18 and

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20 of subarea D, a NNE-SSW striking widely spaced foliation overprints S2 fabric and

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S0, respectively, (Fig. 6; Apx 12) with a moderately to gently NNE plunging

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intersection lineation (L24; L04). Similarly, at location 16 of subarea F (Fig. 9), the same

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foliation crosscuts S0/S1 with steeply NE plunging intersection lineation L0/14. The F4

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folds are open to tight. Although, no clear field evidence was found for S4 overprinting

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S3 in the MB, except for location 16, the relative timing of both fabrics can be inferred

328

from the NNE-SSW strike of D4 that overprints NW-SE striking upright folds (D2/D3)

329

in the NE parts of the Peräpohja Belt (Fig. 1; Lahtinen, unpublished data).

331 332

us

an

M

ed

pt

Ac ce

330

cr

318

3.5. D5 Structures

333

Localized fifth generation mesoscopic F5 folds with N-S and NNE-SSW strike and

334

steeply-dipping axial planes are common throughout the MB. At location 15a of subarea

335

F (Fig. 8), F5 folds are tight, symmetric and associated with a N-S striking evenly

336

spaced S5 cleavage. S5 superposed on F3 folds yields a moderately northward plunging

Page 13 of 49

L35 intersection lineation. At location 9 of subarea E (Fig. 11), F5 folds are closed to

338

tight, parallel in style, generally symmetric and developed in psammitic layers with a

339

sub-vertical intersection lineation (L0/15). Centimeter- to millimeter scale N-S trending

340

folds with a differentiated S5 cleavage can locally be observed in subareas A and B

341

(Fig. 3). Overall, S5 overprints S3 in the eastern part, S2 in the central part and S0/S1 in

342

the western part of the MB. However, direct overprinting relationships between S4 and

343

S5 were not observed. As is the case for S4, the relative timing of D5 structures is

344

deduced from their specific N-S trend and steeply to moderately north plunging

345

intersection lineation in the MB. Moreover, it is difficult to discern S5 from S3 (cf.

346

location 23 of subarea C, Fig. 5, and 15a of subarea F, Fig. 8) or S4 (location 18 of

347

subarea D, Fig. 6), except for the slight differences in their orientation and the

348

intersection lineations (L24, L0/14) associated with S4, which are consistently plunging

349

towards the NE (Fig. 6; Apx12).

cr

us

an

M

ed

350

ip t

337

In summary, the main difference between D5 and D4 related fabric is that S5 is N-S and

352

NNE-SSW striking with steeply to moderately north plunging intersection lineation,

353

whereas D4 is NNE-SSW striking with moderately NE plunging intersection lineation.

354

Similarly, D3 is different than D4 in such a way that S3 is NNW-SSE and WNW-ESE

355

in orientation with pronounced intersection lineations plunging steeply to moderately

356

towards NW, SE and NNW in the MB.

358

Ac ce

357

pt

351

4. Structural interpretation and tectonic events

359 360

An important feature of the high-resolution aeromagnetic image of the MB is that it

361

nicely overlaps with most of the lithological units in the region and hence allowed us to

Page 14 of 49

delineate major structural features (Fig. 2a,b). Despite sparse outcrops, detailed

363

structural analysis of key locations across the MB has revealed multiple near-orthogonal

364

folding events. The near-orthogonal deformation patterns have also been reported from

365

ancient (e.g. Sayab, 2008) and modern tectonic settings (e.g. Shah et al., 2011). The

366

sequence of events, from D1 to D3, are based on fabric overprinting relationship and

367

associated refolded folds, but no direct field evidence has been found to link D3, D4 and

368

D5 other than at locations A2098, A2252 and A1714 in the CLGC (explained below;

369

Fig. 1), where the structural progression from D4 to D5 is noted (Lahtinen, unpublished

370

data). Nevertheless, both D4 and D5 stand out as separate tectonic events and fit well in

371

the regional tectonic framework briefly outlined above and discussed in further detail

372

below. Metamorphic evolution of the MB is deduced from index minerals and is

373

consistent between D2-D3 showing HT-LP greenschist to lower amphibolites facies

374

conditions. Fig. 12 shows area-specific foliation and lineation stereo plots. The same

375

data are also presented as series of rose diagrams to highlight different strike directions

376

for comparison across the MB.

cr

us

an

M

ed

pt

377

ip t

362

The earliest deformation in subareas A and F is defined by steep S1 cleavage with a

379

subvertical L01 intersection lineation. We postulate that S1, and associated F1 parasitic

380

(recumbent) folds, initially formed in response to east-directed low-angle thrusting (Fig.

381

13a), where most of the strain is taken up by black-shales and strain gradient increased

382

with depth towards the basal decollement. Later during the D2 N-S shortening event, F1

383

folds were reoriented about the second generation D2 folds and obtained steep dips of

384

their axial planes, along with S1 and L01 (Fig. 13b). Geometrically, the superimposition

385

of D2 over D1 did not change the orientation of the L1 (AMS) mineral lineation very

386

much. This lineation (well defined by the AMS) generally maintained shallow plunges

Ac ce

378

Page 15 of 49

throughout the effects of D2 except at location 28 where the AMSL is sub-vertical. At

388

this location, centimeter scale well-aligned cordierite porphyroblasts occur, since, the

389

rocks are Mg rich and have low alumina content so that little or no andalusite

390

porphyroblasts formed (Perttunen and Hanski, 2003; Perttunen, 2006). The rocks

391

exposed around subarea E represent relatively deep crustal levels juxtaposed along a

392

fault to lower-grade rocks (Fig. 2a,b). The intensity of D2 deformation continues to

393

decline in the eastern and western parts of the MB where a pervasive S2 is lacking. In

394

the central part of the MB (subareas B, C, D and E), N-S striking S1 is preserved in the

395

hinges of F2 folds or in low-strain pods of D2. The geometry of F1 folds and S1 fabric,

396

both at regional and outcrop scales strongly suggest that prior to D2 an east-vergent

397

tectonic event took place (Lahtinen et al. 2005). The fact that intense D2 deformation is

398

confined to the central portion of the MB is interpreted to reflect a N-S directed

399

shortening gradient with a neck-shape geometry (Fig. 13). Consequently, the D1-D2

400

interference pattern (type 2) produced relatively tight F2 folds in the central part of the

401

MB as compared to the east and west of it (Fig. 13a). The decrease in D2 intensity,

402

especially towards the east, is probably linked to the D2 strain heterogeneity. The

403

mushroom-shaped regional fold pattern in subarea D is interpreted as a F1-F2 fold

404

interference pattern and we further infer that the map pattern in the Mellajoki suite, to

405

the north is due to a D2 antiform, whereas the Martimo suite, to the south is folded into

406

a D2 synform (Fig. 13a).

cr

us

an

M

ed

pt

Ac ce

407

ip t

387

408

D3 deformation is witnessed in the eastern and central part of the MB by the S3

409

cleavage. D3 structure exhibits bimodal orientation with NNW-SSE trends dominating

410

in subarea C, D, and G; and WNW-ESE trends in subarea F. The steep dip of S3

411

suggests that it formed in response to overall NE-SW bulk shortening in the region

Page 16 of 49

superimposed on D2 and D1 structures. Syn-tectonic cordierite porphyroblasts aligned

413

along S3 suggest at least lower amphibolites facies metamorphism until the end of D3.

414

Similarly, S4 can be assumed to have formed due to NW-SE shortening (D4) producing

415

a steeply dipping foliation. Although D3-D4 overprinting relationships were not found

416

in the field, however, the relative timing of these events can be extrapolated from the

417

relative timing between D4 and D5 in the CLGC (Lahtinen, unpublished data), and

418

between D2 and D3 in the central part of the MB. Thus, we infer that D4 is a separate

419

tectonic event caused by NW-SE crustal shortening at ca. 1.83-1.81 Ga (Fig. 14). The

420

last recognized deformation event (D5) affected almost the entire MB and is

421

characterized by broadly WSW-ENE or E-W oriented crustal shortening. It is more

422

pronounced in the eastern part of the MB whose regional oval-shape geometry can be

423

attributed to the D5 event (Fig. 13c). Multiple sets of intersection lineations ((L01, L12,

424

L0/15, L23, L24, L35, L25; terminology after Bell and Duncan, 1978) and their associated

425

S1, S2, S3, S4 and S5 foliations bear witness of multiple near-orthogonal shortening

426

events.

cr

us

an

M

ed

pt

427

ip t

412

A reconnaissance study in the CLGC (Bedrock of Finland – DigiKP; Lahtinen,

429

unpublished data) indicates that the youngest ductile deformation event (D5 of this

430

study) produced subhorizontal shearing and an associated mineral lineation with

431

vergence towards the NE (Fig. 1). The Lohiniva appinite suite plutonic rocks are about

432

1.80 Ga in age and crosscut earlier structures but they also show in some cases this NE-

433

SW oriented subhorisontal mineral lineation and shearing (e.g., A1714 in Fig. 1). The

434

NE vergent thrusting (ca. 1.80 Ga) observed in the Hanhimaa-Rautuvaara area (H in

435

Fig. 1; Niiranen et al., 2007) and reverse faulting with NE vergence at locations 10 and

436

11 (Fig. 2a) are linked with the D5 event of this study. At location A2098 (Fig. 1),

Ac ce

428

Page 17 of 49

strong mineral stretching seen as rodding and sheath folds, moderately dipping towards

438

SW, overprint vertically turned NE-SW oriented migmatites with 1.86-1.85 Ga

439

metamorphic zircon overgrowths (Lahtinen et al., 2015). Similarly, a strongly deformed

440

granite A2252 (< 1.87 Ga; Fig. 1) has been vertically turned with NE-SW axial trace.

441

Also east of location A1714 in Fig. 1 there are observations where migmatites are

442

uprightly folded with NE-SW axial trace inclined to SE and later overprinted by

443

subhorisontal shearing and mineral lineation with NE-SW orientation. Thus, we

444

designate this folding with NE-SW axial trace as D4 overprinted by D5 with age ranges

445

of 1.86/1.85-1.81 Ga and 1.81-1.77 Ga, respectively. The lower age limit of D5 is based

446

on the strong 1.79-1.77 Ga metamorphic event recorded on titanite and monazite ages in

447

wide area around the boundary zone in Fig. 1 (Hiltunen, 1982; Lehtonen, 1984;

448

Mänttäri, 1995; Rastas et al., 2001; Perttunen and Vaasjoki, 2001; Väänänen and

449

Lehtonen, 2001; Bergman et al., 2006; Niiranen et al., 2007; Lahtinen et al., 2015). This

450

age period seems to have been the last major thermal event before the stabilization of

451

the crust in the study area.

454

cr

us

an

M

ed

pt

453

5. Regional tectonic implications

Ac ce

452

ip t

437

455

Based on structural field observations, aeromagnetic interpretations, AMS data and

456

regional reconnaissance survey, we propose a schematic model (Fig. 14) to explain the

457

Paleoproterozoic tectonic evolution of Northern Fennoscandia between 1.92- 1.77 Ga.

458

The rifting of the Archean continent (pre-Karelia) initiated with a large scale plume

459

event at 2.1 Ga (Vuollo and Huhma, 2005; Hanski and Huhma, 2005, Lahtinen et al.,

460

2015b), followed by triple point-type continental breakup forming NNW-SSE trending

Page 18 of 49

461

(present coordinates) passive margin and a ENE-WSW trending aulacogen, now

462

represented by the CLGC (Fig. 1; Lahtinen et al., 2015a).

463 The onset of collisional stage is seen as an east vergent thrusting where both the Kittilä

465

allochthon and the Martimo belt (D1, this study) show thin-skin style thrusting. The D1-

466

D2 in the Martimo belt formed between 1.91-1.88 Ga based the maximum

467

sedimentation age of the younger Martimo sediments at ca. 1.91 Ga (Ranta et al., 2015)

468

and D2 crosscutting plutons with zircon U-Pb ages between 1886±10 and 1880±2 Ma

469

(Perttunen and Vaasjoki, 2001). Diorite A490 (Fig. 1) had also a concordant monazite

470

registering an age of 1895±3 Ma (ibid.), which either indicates older intrusion age

471

(zircon from cooling stage) or that the monazites are xenocrysts from D1 metamorphic

472

peak. The exact timing of D1 in the MB is still open but an age older than 1.89 Ga is

473

postulated. The timing of the D1 event in the Kittilä area is ≥1.91 Ga based on the

474

crosscutting pluton (A1206 in Fig. 1), causing contact metamorphic aureole (Rastas et

475

al., 2001). The graphite- and Fe-sulfide-bearing Liekopalo paraschists and Kaskimaa

476

turbidites of the Martimo belt are considered as passive margin and/or fore-deep

477

sedimentary rocks with detritus from the Karelia lower plate, whereas the deposition of

478

the Ristivuoma sedimentary rocks is coincident with the formation of foreland fold-and-

479

thrust belt with a large contribution of detritus from the eroding Paleoproterozoic arc

480

rocks in the west (Lahtinen et al., 2015a).

cr

us

an

M

ed

pt

Ac ce

481

ip t

464

482

The Mellajoki suite are more competent, due to abundant quartzites and granites than

483

the Martimo belt rocks and show a relatively more thick-skin type thrusting (Fig. 2a),

484

possibly including the Archean basement involvement (thick-skin in sensu stricto) close

485

to the western margin of the Karelia continent. The Peräpohja Belt proper (Fig. 1) is

Page 19 of 49

characterized by competent quartzites. Anyhow, preliminary results indicate that early

487

E-vergent shortening has affected also this belt (Sayab, unpublished results). Coincident

488

with the proposed collision between Norrbotten and Karelia the Lapland-Kola orogen

489

formed during a continent-continent collision (Fig. 14) between the Karelia and Kola

490

continents (Lahtinen et al. 2005; Daly et al. 2006 and references therein).

ip t

486

491

The continued collision between the Keitele continental ribbon and Savo arc with the

493

Karelia continent at 1.90-1.89 Ga (Fig. 14) was associated with major elongation of

494

structures (L1 and F1 coincides, i.e., sheath folds) turning towards NE in the

495

Outokumpu area (Koistinen, 1981). We interpret that due to this NE push the large

496

Pudasjärvi Archean lithosphere moved almost orogen parallel towards N and started a

497

partial basin inversion in the aulacogen. This shortening is mainly seen as upright

498

folding (D2) and faulting characterized in the Peräpohja and Martimo belts. The folding

499

varies from typical open to tight at the junction of Martimo and Mellajoki suites (areas

500

B and C in Fig. 2a).

us

an

M

ed

pt

501

cr

492

D3 crosscuts D2 and thus, has to be younger than 1.88 Ga, based on 1.89-1.88 Ga D2

503

plutons, and older than 1.86 Ga, based on metamorphic peak at 1.86-1.85 Ga (A2098 in

504

Fig. 1; Lahtinen et al., 2015a), but the exact timing is to be determined. Here we follow

505

the model of Lahtinen et al. (2014) and infer that a major accretion and collision stage

506

between 1.88-1.87 Ga (D3) is responsible for strong shortening in the central part of

507

Fennoscandia followed by 1.87-1.86 Ga buckling forming oroclines. The D3 shortening

508

is prominent in the eastern MB where the earlier (D2) N-S shortening is seen only in

509

open upright folding without axial cleavage. The same holds for the western MB, but

510

there younger D3 effects are less pronounced and localized. We propose that the eastern

Ac ce

502

Page 20 of 49

MB reflects a strain shadow formed during N-S shortening of NW-SE trending horst

512

and graben structures with transform faults, which implies a strong Archean basement

513

involvement. Moreover, it is tentatively proposed that the NE directed thrusting close to

514

the cryptic suture (Fig. 2b) and NW-SE trending belts east of our study area (Fig. 1) are

515

formed due to the D3 event. The S3 foliation is generally NNW-SSE striking, however,

516

WNW-ESE trends are not uncommon, and it is possible that the D3 could be separated

517

to collision- and buckling-related stages but more data is needed to verify this.

cr

ip t

511

us

518

In Fig. 14 we only show major shortening stages but the abundant occurrence of 1.88-

520

1.87 Ga stitching plutons in a 100-150 km wide zone on the cryptic suture (Fig. 1)

521

favors extensional and/or transtensional stages during or after the D2-D3 events. The ca.

522

1.85 Ga metamorphic peak with associated magmatism in Northern Fennoscandia

523

(Bergman et al., 2006; Lahtinen et al., 2015a) is tentatively linked with D3 extensional

524

(collapse) event occurring after the proposed buckling (Fig. 14). We correlate this to the

525

rifting period at 1.85-1.83 Ga documented in southern Finland (Lahtinen and Nironen,

526

2010; Nironen and Mänttäri, 2012).

M

ed

pt

Ac ce

527

an

519

528

D4 is not well established but upright to SE inclined folding with NE-SW axial traces

529

are found in the CLGC and also in NE part of the Peräpohja belt. Following lines of

530

evidence support D4 deformation, 1) they fold ca.1.85 Ga migmatites and ≤1.87 Ga

531

granites (Fig. 1), 2) NE-SW trending folds in subareas D and F, 3) NE-SW trending

532

lineaments on the aeromagnetic image and 4) NE-SW oriented SE-side up reverse faults

533

between the Rovaniemi and Mellajoki suites and within the Mellajoki suite (Fig. 2b).

534

We link this event to strong shortening occurred in southern Finland at ca. 1.83-1.82 Ga

535

(Ehlers et al., 1993; Väisänen and Hölttä, 1999; Pajunen et al., 2008; Skyttä and

Page 21 of 49

536

Mänttäri, 2008; Saalmann et al., 2009; Nironen and Mänttäri, 2012). Lahtinen et al.

537

(2015a) tentatively proposed that 1.82– 1.78 Ga metamorphic ages in Northern

538

Fennoscandia can possibly be separated into two events at 1.83–1.82 Ga and 1.79–1.77

539

Ga, possibly linked to D4 and D5, respectively.

ip t

540 Small 1.80 Ga (e.g., A1714 in Fig. 1) appinitic plutons and dykes are common in the

542

CLGC and pegmatitic to aplitic granites with 1.80-1.77 Ga ages are abundant in the

543

CLGC and the Pajala shear zone (e.g. Bergman et al., 2001; Ahtonen et al., 2007;

544

Åkerman and Kero, 2012; Lauri et al., 2012). We infer that the 1.80 Ga bimodal suite is

545

related to a D4 extension, which has partly been a collapse-type event. The D5

546

deformation as a whole is seen as a strong 1.79-.77 Ga tectono-metamorphic event in

547

Northern Fennoscandia (Mänttäri, 1995; Rastas et al., 2001; Perttunen and Vaasjoki,

548

2001; Väänänen and Lehtonen, 2001; Corfu and Evins, 2002; Bergman et al., 2001,

549

2006; Niiranen et al., 2007, Lahtinen et al., 2015a). We interpret that during D5 event,

550

(1. in Fig. 14) open folds with approximately N-S (NNW-SSE) trending axial planes

551

formed due to orthogonal collision between two rigid crustal units, Norrbotten and

552

Karelia. A change towards more transpressive compression produced a strong stretching

553

with vergence towards NE (2. in Fig 14) expressed as subhorizontal shear planes and

554

mineral lineation in the aulacogen (CLGC) and as more steep oblique reverse faulting

555

with dextral-strike slip component along the cryptic suture (Fig. 1). Locally, reverse

556

faults along more competent units, like the Mellajoki suite, produced closed folds with

557

steep fold axis, e.g., area E in Figs. 2a,b and 13. We attribute D5 to a major continent-

558

continent collision at the western part of Fennoscandia (Fig. 14) following the tectonic

559

evolution proposed by Bergh et al. (2010) for the West Troms Basement Complex in

560

the northwestern part of Fennoscandia. They model a continued NE-SW orthogonal

Ac ce

pt

ed

M

an

us

cr

541

Page 22 of 49

561

shortening with increasing transpressive component (see also Larsen Angvik, 2014) at

562

1.80-1.76 Ga which includes an early NE-directed thrusting and SW-dipping crustal

563

detachment/foliation in medium to high-grade metamorphic conditions. They also

564

proposed late orogen parallel local SE-directed thrusting.

ip t

565 The model in Fig. 14 depicts only the period between 1.92-1.77 Ga of the

567

Paleoproterozoic Fennoscadian tectonic evolution. The late component in the Pajala

568

shear zone is a regional sinistral sense of shear (Wikström et al. 1996) with local

569

eastern-side-up movements in the western branch (Bergman et al., 2001). Clear brittle-

570

type NW-SE oriented faulting (Fig. 2b) in the CLGC is also evident from aeromagnetic

571

images but the age of this faulting is unknown. As a whole strain heterogeneity is a very

572

important feature in Northern Fennoscandia and in local studies important events can be

573

left unnoticed or the observed shortening directions do not correlate with the large scale

574

plate movements. An important feature is the occurrence of rigid crustal blocks, pieces

575

of Archean lithosphere, which can move orogen parallel (e.g, in D2) or first produce

576

orthogonal structures followed by transpression (e.g. D5). Smaller rigid to more

577

competent units like horst and graben blocks can form sheltered domains, escaping

578

shortening (e.g., D2 followed by D3). One major feature is the aulacogen, which has

579

been thickened by thrusting during orthogonal or nearly orthogonal shortening in D1,

580

D3 and D5, and shows basin inversion during D2 and D4.

us

an

M

ed

pt

Ac ce

581

cr

566

582

The sequence of tectonic events described herein (D1/D2, D3, D4 and D5) are

583

correlated with the formation of Lapland-Savo, Fennian, Svecobaltic and Nordic

584

orogens of the composite Svecofennian orogen (1.92–1.79 Ga; Lahtinen et al, 2005;

585

2009; 2014), respectively. These orogens are defined based on the information received

Page 23 of 49

from the Fennoscandian Shield and their continuation towards present north is

587

unknown. D2 has been linked to orogen parallel shortening but another or coincident

588

tectonic event could have been a collision in the north. Similarly, D4 could have been

589

caused by simultaneous collisions from NNW and SSE and buckling indicates plate

590

movement in N-S direction. Greenland has an active Paleoproterozoic tectonic history

591

between 1.88 and 1.80 Ga (e.g. Garde and Hollis, 2010), which makes it possible

592

candidate for this "terra ingcognita". We prefer the buckling mechanism for generating

593

the oroclines (Lahtinen et al., 2014) but a subduction rollback with bending in a strike-

594

slip setting (Rosenbaum et al., 2012) can be another possibility for orocline formation.

595

The latter model would fit to the proposed accretionary crustal growth towards

596

southwest during and after the Fennian Orogeny (Hermannson et al., 2008; Saalman et

597

al., 2009; Bogdanova et al., 2015). However, in both scenarios the bi-model

598

distribution of D3 fabric suggests that D3 initially formed due to the NE-SW

599

shortening.

cr

us

an

M

ed

600

ip t

586

Several orthogonal tectonic events in northern Fennoscandia indicate amalgamation of

602

crustal blocks around Archean Karelia continent core at 1.92-1.77 Ga. Fennoscandia

603

was then integrated as a key component in the Columbia/Hudsonia/Nuna supercontinent

604

(Zhao et al., 2002, 2004). One stage of the supercontinent formation can be seen in an

605

obvious correlation between the c. 1.8 Ga Ketilidian (Greenland) and Nordic orogens

606

both characterized by linear Andean-type batholiths of similar age (Transcandinavian in

607

Fig. 14). The Nordic orogen is interpreted as continent-continent collision but the

608

Ketilidian orogen displays a convergent plate-tectonic system without subsequent

609

collision (Garde et al., 2002). If the latter interpretation holds for Fennoscandia the D5

610

tectonic event and the Nordic orogen would be related to advancing accretionary orogen

Ac ce

pt

601

Page 24 of 49

611

(Andean-type) with retro-arc fold and thrust belts (Lahtinen et al., 2009; Saalman et al.,

612

2009; Bogdanova et al., 2015). Comprehensive supercontinent correlations are outside

613

of the scope of this study but it is evident that our results are important in those

614

considerations.

616

ip t

615 6. Conclusions

cr

617

1) Based on fabric overprinting relationship and associated refolded folds, we have

619

three deformation events (D1 to D3), which can be recognized in the study area.

620

Although no direct field relationships have been found to link D3, D4 and D5

621

structures, however, age and field data outside the Martimo strongly suggest D4 and D5

622

are separate younger tectonic events.

M

an

us

618

623

2) We interpret that D1 structures initially formed in response to east directed low-angle

625

thrusting, and during a subsequent N-S shortening event (D2) were reoriented about F2

626

folds. Indeed, the subvertical nature of F1 fold axis (F01) is due to the later effects of D2

627

shortening.

pt

Ac ce

628

ed

624

629

3) D2 produced upright folds within a major synform (MB)- antiform (Mellajoki suite)

630

pair, where the variation in the D2 strain produced a neck-shape geometry at the centre

631

of the MB due to more intense D2 deformation in this central portion, whereas in the

632

eastern and western parts a pervasive S2 is lacking.

633 634

4) Map-scale D3 deformation is pronounced in the eastern part, whereas it is more

635

localized towards the western part. The orientation of S3 fabric suggests that it formed

Page 25 of 49

636

in response to NE-SW bulk shortening in the region superimposed on D2 and D1

637

structures.

638 5) D4, formed due to NW-SE shortening, is less pronounce in the study area and is seen

640

locally as a steeply dipping foliation (S4) and or faulting along the contacts of

641

lithological units. The last recognized deformation event (D5), characterized by broadly

642

WSW-ENE oriented crustal shortening, is prominent in the eastern part of the MB

643

whose regional oval-shape geometry can be attributed to the D5 event.

us

cr

ip t

639

644

6) Multiple sets of intersection lineations (L01, L12, L0/15, L23, L24, L35, L25) and their

646

associated S1, S2, S3, S4 and S5 foliations suggest multiple near-orthogonal shortening

647

events. Strain heterogeneity appears to be responsible for the lack of individual post-D1

648

deformation event in specific areas. We propose that the Archean basement geometry

649

with horst and graben structures and transform faults have played important role in

650

producing variations in strain gradient.

ed

M

an

645

pt

651

7) Our D1 to D5 deformation scheme supports the formation and tectonic evolution of

653

the Lapland-Savo, Fennian, Svecobaltic and Nordic orogens of the composite

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Svecofennian orogen (1.92–1.79 Ga; Lahtinen et al, 2005; 2009; 2014).

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Acknowledgements

We would like to give gratitude for the mapping geologists of 70's, especially Vesa

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Perttunen and Jukka Väänänen, who's mapping results form the basis of our work. The

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first author would like to thank Tuomo Manninen with whom the project in this area

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was initiated. Fruitful discussions with Pentti Hölttä, Mikko Nironen and Laura Lauri

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are also acknowledged. Pietari Skyttä is acknowledged for useful discussions in the

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field.

665 References

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Figure captions

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Fig. 1. Simplified geological map of the study area modified from the Bedrock of Finland – DigiKP and SGU bedrock map of Sweden in 1:1 million scale. Stars and boxes refer to samples with zircon ages discussed in the text. Data from Huhma (1986), Rastas et al. (2001), Perttunen and Vaasjoki (2001), Väänänen and Lehtonen (2001), Ahtonen et al.(2007), Lahtinen et al. (2015a) and Lahtinen (unpublished data). CLGB – Central Lapland Greenstone Belt; H – Hannukainen-Rautuvaara area; P –Pajala town. Inset figure shows the location of the study area in Fennoscandia.

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Fig. 2a. Simplified geological map of the study area modified from Bedrock of FinlandDigiKP with subareas and observation points. For the Norrbotten Province see legend in Fig. 1.

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Fig. 2b. Aeromagnetic map (Korhonen et al., 2002) of the study area (Fig. 2a) with main lithological boundaries and axial traces of deformation (Fig. 2a).

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Fig. 3a,b. Field photograph and line drawing of F1 folds with steeply dipping S1 cleavage. L01 is sub-vertical. S5 cleavage is N-S striking and overprints S0/S1. AMS mineral lineation is shallowly plunging towards the west. c) Stereographic plots of foliations and lineations. Rose diagram showing the relationship between S1 and S5 orientation. Location 3, subarea A.

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Fig. 4 (a-c) Field photograph and line diagram showing overprinting relationships of S0/S1, S2 and S5. The N-S orientation of S0/S1 is obtained from the hinge of F2 folds. S5 is N-S striking and overprints S0/S1 and S2. AMS mineral lineation is gently plunging towards east. d) Poles (black dots) and moving average rose diagrams (dark blue) for S0/S1, S2 and S5 foliations, and L1 mineral lineation deduced by AMS (AMSL1). Location 8, subarea B.

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Fig. 5. (a) A unique 3-D outcrop location showing overprinting relationship of S0/S1, S2 and S3 and L01, L12 and L23 intersection lineations. F3-F2 fold interference yields type-1 geometry. (b) Showing relationship of intersection lineations (L01, L12 and L23) and moving average rose diagram for S1, S2 and S3. Location 23, subarea C. Fig. 6. Photograph and sketch showing overprinting effects of D4 over S2. S4 cleavage is NE-SW striking and steeply dipping with gently NE plunging L24 intersection lineation. Rose diagram showing poles and strike orientation of S2 and S4. Location 18, subarea D. Fig. 7. (a,b) Outcrop showing the relationship between S0/S1 and S3 foliations. Cordierite porphyroblasts along the S0/S1 are coarser in size than the ones developed along S3. S3 foliation is NNW-SSE and clearly crosscuts S0/S1. AMS mineral 745 lineation is steeply plunging at this location (see text for further description). Location 28, subarea E.

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Fig. 8. (a) Field photos and sketches showing overprinting relationships of L0/13 and L0/1/35 and associated S0/S1, S3 and S5 fabrics. The N-S striking S0/S1 is crenulated by WNW-ESE striking S3. D5 produced tight N-S trending folds with moderately N plunging intersection lineation. (b) Close-up showing the orientation of S5 and associated intersection lineation (L0/1/35). (c) Poles to foliation planes (black dots) and moving average rose diagrams for S0/S1, S3 and S5 and associated intersection lineations (L0/13 , L0/1/35. Location 15a, subarea F.

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Fig. 9. (a) Field sketch and photos showing tight N-S trending F1 folds (in the east), which are symmetrically buckled by the D3. F3 folds are open with penetrative WNWESE striking S3 axial plane cleavage. F4 folds (in the west) are open to tight, where the S4 is NE-SW striking and steeply dipping. F1 and F4 folds can be distinguished based on the fact that F1 folds are buckled due to the effects of D3, whereas F4 folds are the youngest in this location. (b) Poles (black dots) and moving average rose diagrams for S0/S1, S1, S3 and S4 and associated intersection lineations (L01, L0/14, L0/13). Location 16, subarea F.

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Fig. 10. Outcrop showing the effects of S3 on S0/S1. S3 clearly overprints on S0/S1. Location 30, subarea G.

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Fig. 11. N-S trending F5 fold with steeply plunging intersection lineation (L0/1 5). Note, that it is difficult to discern F5 folds from F3 (cf. Fig. 5) or F4 (cf. Fig. 6), except for the slight differences in their trends and the L2 4 intersection lineation associated with the S4, which is moderately plunging towards the NE. Location 9, subarea E.

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Fig. 12. Area-specific lower hemisphere equal-area stereographic plots of foliations and lineations measured in the field to establish fabric overprinting relationships. S5 is generally N-S striking (subareas A,B,E,F), S4 is NW-SE striking (subareas D,F), S3 has both NNW-SSE and WNW-ESE (subareas C-G) orientation due to the later buckling effect, and S2 is E-W striking (subareas A,B,C,D) with local variations. S0/S1 is reoriented due to the effects of D2. Fig. 13. Tectonic successions and kinematic model for the Martimo metasedimentary belt and Mellajoki suite. D1 to D5 tectonic events are modeled based on fabric crosscutting relationship (D1-D3) and specific orientations (D4 and D5). We infer that D1, D2, D3 and D5 were the major tectonic events shaping and reshaping the Martimo belt, whereas D4 was relatively weak in the sequence.

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Fig. 14. Sketch illustrating the proposed tectonic model for Northern Fennoscandia at 1.92–1.77 Ga. D1- D5 refer to main shortening events described in this study. P Pudasjärvi crustal block; O - Outokumpua area. Note that Paleoproterozoic arc turns to Paleoproterozoic crust after collision/accretion. Arc and microcontinent names refer to Lahtinen et al. (2005; 2014).

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