Syn-extensional faulting controlling structural inversion – Insights from the Palaeoproterozoic Vargfors syncline, Skellefte mining district, Sweden

Precambrian Research 191 (2011) 166–183

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Syn-extensional faulting controlling structural inversion – Insights from the Palaeoproterozoic Vargfors syncline, Skellefte mining district, Sweden Tobias E. Bauer a,∗ , Pietari Skyttä a , Rodney L. Allen a,b , Pär Weihed a a b

Division of Geosciences and Environmental Engineering, Luleå University of Technology, SE-971 87 Luleå, Sweden Boliden Mineral AB, SE-936 81 Boliden, Sweden

a r t i c l e

i n f o

Article history: Received 26 October 2010 Received in revised form 25 September 2011 Accepted 27 September 2011 Available online 5 October 2011 Keywords: Palaeoproterozoic Skellefte district 2D forward modelling Structural analysis Basin inversion VMS deposits

a b s t r a c t The Vargfors basin in the central Skellefte district, Sweden, is an inverted sedimentary sub-basin within a Palaeoproterozoic (1.89 Ga) marine volcanic arc. The sub-basin formed from upper-crustal extension and subsequent compression, following a period of intense marine volcanism and VMS ore formation. Detailed mapping and structural analysis reveals a pattern of SE–NW-striking normal faults and interlinked NE–SW-striking transfer faults, which define distinct fault-bound compartments, each with an individual structural geometry and stratigraphy. Constraints on the deformation style and mechanisms achieved by 2D forward modelling are in agreement with the previously inferred inversion of the early normal faults during a regional crustal shortening event. A rheologically weak carbonate-rich layer at the base of the sedimentary sequence favoured the fault inversion over more distributed shortening as the controlling deformation mechanism. Transposition of sedimentary strata into the approximately SE–NW faults led to formation of asymmetric synclines that were tightened during progressive shortening. Structural analysis infers a progressive opening of the basin towards SE and NW with time. Furthermore, it is inferred that a displacement gradient was developed along the main structural grain, with decreasing dip-slip displacements towards SE and NW, both during the extension and the structural inversion.VMS deposits in the vicinity of the contact between the volcanic and the overlying sedimentary rocks were formed along early normal faults, which reacted as fluid conduits. Subsequently, the deposits were transposed into the inverted faults during crustal shortening. Consequently, the inverted faults provide a useful tool for mineral exploration in the district. © 2011 Elsevier B.V. All rights reserved.

1. Introduction The Skellefte district in northern Sweden (Fig. 1) covers a 120 by 30 km area of Palaeoproterozoic volcanic, intrusive and sedimentary rocks and is one of the most important mining districts in Europe, producing Zn, Cu, Pb, Ag and Au from volcanogenic massive sulphide (VMS) and orogenic gold deposits. Allen et al. (1996) inferred that a complex horst-and-graben geometry was developed during crustal extension in the Skellefte district and, consequently pointed out the importance of the early extensional faults influencing the evolution of the volcanic and sedimentary successions in the district. Massive sulphide deposits occur preferentially in the upper part of the main volcanic succession (Skellefte Group), close to the contact with an overlying mainly sedimentary succession (Vargfors Group; Allen et al., 1996). However, few constraints have been provided about the geometry of the

∗ Corresponding author. Tel.: +46 76 8049272. E-mail address: [email protected] (T.E. Bauer). 0301-9268/$ – see front matter © 2011 Elsevier B.V. All rights reserved. doi:10.1016/j.precamres.2011.09.014

fault pattern, the contact relationships between the Skellefte Group and the Vargfors Group, or how, if at all, these features affected the subsequent deformation. Moreover, few previous studies focusing in the crustal evolution of the area have assessed the coupling between stratigraphy and structure which is the main approach of this paper. As the main outcome we describe the complex evolution of a sedimentary sub-basin within an active volcanic arc at the change from crustal extension to compression which might be further associated with the accretion to the Archean continent (Karelian craton margin; Weihed et al., 2002a, 2005; Lahtinen et al., 2004, 2005). Rocks in the Skellefte district are deformed and metamorphosed to greenschist or amphibolite facies. However, the Vargfors syncline shows relatively low strain, the lowest metamorphic grade and therefore the best visible primary features in the whole Skellefte district, hence making it the most suitable area to study the evolution of the stratigraphy and structure. The results of this case study provide insights into the structural style of and strain partitioning within an inverted Palaeoproterozoic basin, thus also setting the framework for future interpretations on the structural

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Fig. 1. Simplified geological map of the Skellefte district and surroundings. Modified after Kathol et al. (2005). Location of the central Skellefte district marked on the map.

geometry and evolution of the Skellefte mining district. Furthermore, some insights into localisation and subsequent tectonic transposition of the VMS deposits are provided. Results from mapping of outcrops, drill cores and thin sections are integrated into a detailed geological map and several cross-sections elucidating the present-day lithostratigraphy and structure of the central part of the Vargfors syncline. 2D-foreward modelling was applied for cross-sections transecting selected compartments within the Vargfors syncline to constrain the succession of deformational events as well as to evaluate the controlling deformation mechanism in forming the syncline. Based on these results a tentative model for each of the fault-bound compartments is presented. These data are further synthesized to provide a kinematic and geometric framework for the inverted Vargfors syncline and a three-dimensional model showing the restored geometry for it. 2. Geological background The Skellefte district comprises 1.90–1.86 Ga, Palaeoproterozoic metamorphosed supracrustal and associated intrusive rocks occurring as a 120 by 30 km area along the Skellefte river (Fig. 1). Although all rocks in the study area were metamorphosed under

greenschist conditions, we use primary sedimentary and volcanic terminology to emphasize the mostly well preserved primary lithological textures. The 1.90–1.88 Ga Skellefte Group is the lowest stratigraphic unit in the central Skellefte district and comprises mainly subaqueous lava domes, porphyritic cryptodomes, lavas and volcaniclastic rocks with largely rhyolitic and minor basaltic, andesitic and dacitic composition. Common sedimentary intercalations include grey to black mudstone, volcaniclastic siltstone, sandstone, breccia-conglomerate, volcaniclastic rocks with carbonate matrix and rare limestone (Allen et al., 1996; Kathol and Weihed, 2005; Montelius et al., 2007). As the base of the Skellefte Group is not exposed, the stratigraphic thickness is not known, but exceeds 3 km in the N-part of the central Skellefte district (Allen et al., 1996). The Vargfors Group overlies the Skellefte Group, and comprises argillites, sandstones, conglomerates and subordinate carbonaterich mudstones to conglomerates, occurring throughout the district. In the central part of the Skellefte district, Vargfors Group sedimentary rocks occur in a sedimentary sub-basin, the so called Vargfors basin or Vargfors syncline (Figs. 2 and 3; c.f. Kautsky, 1957; Dumas, 1986; Kathol and Weihed, 2005) where a pronounced pattern of NW–SE-striking faults crosscut by NE–SW-striking faults

168 T.E. Bauer et al. / Precambrian Research 191 (2011) 166–183 Fig. 2. Lithological map of the central Vargfors basin. Including lithological and tectonic contact relationships as well as observation points. Modified after Allen et al. (1996). Bold roman numbers (I–VI) refer to individual structural compartments within the basin. LS W: long section west; LS E: long section east. Coordinates given in the Swedish national RT-90 grid.

T.E. Bauer et al. / Precambrian Research 191 (2011) 166–183 Fig. 3. Structural map of the central Vargfors basin. Modified after Allen et al. (1996), based on 642 structural measurements taken during this investigation from 489 localities (see Fig. 2 for location of observation points). Bold roman numbers (I–VI) refer to individual structural compartments within the basin. Coordinates given in the Swedish national RT-90 grid.

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define specific fault-bound compartments (Allen et al., 1996; Bauer et al., 2009). The character of the Skellefte–Vargfors Group contact varies from one compartment to the next, being either conformable, disconformable, or faulted at different locations (Allen et al., 1996). Minor disconformities occur also within the sedimentary package. The uppermost part of the Vargfors Group consists of basalts and andesites related to the Gallejaur complex (1873 ± 10 Ma; Skiöld, 1988). To the N of the Vargfors syncline, the central Skellefte district supracrustal rocks are bound by the Jörn intrusive complex which comprises four main intrusive phases (GI–GIV; 1.89–1.86 Ga; Wilson et al., 1987; Gonzáles Roldán, 2010) ranging from gabbro to granite, with tonalite and granodiorite most abundant (Fig. 1). Tonalite clasts derived from the oldest intrusive phase (GI) occur within the Vargfors Group conglomerates and indicate that there was only a few million years between the emplacement of the Jörn intrusive complex, and its subsequent uplift and erosion. The Smargin of the central Skellefte district is bordered by an extensive region of metasedimentary rocks (Bothnian Basin) and granites to syenites ranging from pre- to post-Skellefte Group in age. To the NW, the central Skellefte district has a poorly defined boundary with the Arvidsjaur Group (1876 ± 3 Ma; Skiöld et al., 1993), dominated by subaerial felsic volcanic rocks and minor sedimentary rocks. Pre-1.88 Ga regional deformation has been suggested for the Skellefte district and its surroundings, evident from undeformed approximately 1.87 Ga rocks intruding into tectonically foliated approximately 1.89 Ga Jörn GI-type granitoids (Lundström and Antal, 2000). Rutland et al. (2001a,b) and Skiöld and Rutland (2006) suggested an age older than 1.89 Ga for their earliest tectonometamorphic event related to crustal extension (their D1 ). However, this interpretation is controversial due to structural and intrusive relationships within the Bothnian Basin and the Skellefte district (Weihed, 2003) and due to a critical evaluation of the geochronological studies (Högdahl and Lundqvist, 2009). A layer-parallel foliation occurring throughout the district, mainly in clastic volcanic and sedimentary rocks, but also locally in competent lavas and intrusions is the oldest fabric in the district (Allen et al., 1996). No folds related to this foliation have been found which led Allen et al. (1996) to interpret the fabric partly due to diagenetic compaction and partly as an early tectonic fabric. The main compressional deformation within the Skellefte district (at 1.87–1.82 Ga) is suggested to relate either to dextral transpression arising from SE–NW crustal shortening (Bergman Weihed, 2001) or, alternatively, to SW–NE shortening (Weihed et al., 2002a). Characteristic structures developed during this event are NW-striking faults, typically with S-side-up kinematics and gently plunging upright folds (Bergman Weihed, 2001). Allen et al. (1996) related the variations in the plunge direction of the fold axes with steep N-striking, mainly dipslip cross faults and N-striking folds, consequently attributed to a later phase of deformation. This deformation corresponds to the D3 of Bergman Weihed (2001), which involved deformation along steep N-S-striking major shear zones and crenulations of older deformation fabrics, both reflection approximately E-W crustal shortening (Bergman Weihed, 2001). The post 1.88 Ga age for the main folding in the central Skellefte district is supported by the similar trends and styles of deformation structures both within the Vargfors and Skellefte Group rocks (Allen et al., 1996) and by the 1865–1850 Ma age for the D2 deformation obtained by Skiöld and Rutland (2006) to the S of the Skellefte district. The latest major deformation event took place at approximately 1.80 Ga (Bergman Weihed, 2001). Metamorphism in the E-part of the Skellefte district reached its peak during the oldest deformation event prior to the ∼1875 Ma intrusions (Lundström et al., 1997, 1999). In the central part of the district the metamorphic peak is inferred at approximately

1825 Ma (Weihed et al., 2002a). Within the central Skellefte district, metamorphic conditions varied from greenschist facies in the N (∼400 ◦ C and ∼2 kbar) to amphibolite facies in the S (Kathol and Weihed, 2005). The latest phase of intrusions are the late- to posttectonic 1.82–1.78 Ga Revsund-type granites that form a part of the Transscandinavian Igneous Belt (TIB; Kathol and Weihed, 2005). 3. Lithologies in the Vargfors syncline Successions of sandstone turbidites, interbedded grey mudstones, and monomict conglomerates with unaltered and hydrothermally altered clasts from the surrounding Skellefte Group volcanic rocks dominate the central part of the Vargfors syncline (Fig. 2; Dumas, 1986; this study). Variations in the clast composition and degree of rounding of clasts within the less common polymict conglomerates reflect different sedimentary sources (Table 1). In addition, some rocks, especially along the main high-strain zones, have a significant calcite component. Secondary calcite, observed both as pseudomorphic replacement of plagioclase and calcite veining (Fig. 4a) is common in nearly all the sedimentary rocks within the Vargfors syncline, but is most abundant within the high-strain zones. All the sedimentary rocks are intruded by felsic and mafic, post-tectonic sills and dykes. 3.1. Carbonate-rich sedimentary rocks Fine-grained, grey, calcite-rich mudstones occur along the Scontact of compartments V and VI (Fig. 2). They show traces of transposed bedding seen as alternating layers of calcite and quartz grains (Table 1 and Fig. 4b). Lenses or patches of carbonatecemented conglomerate occur within compartments II–VI (Fig. 2). This conglomerate comprises mainly felsic volcanic clasts embedded in a calcite-rich matrix (Fig. 4c and d). Intense weathering of the carbonate matrix has imparted a characteristic dark brown-black colour to the outcrops. 3.2. Turbiditic units These units crop out throughout the study area, and comprise alternating layers of fine-laminated mudstone, greywacke, arkose, and local interbeds of conglomerate (Fig. 4e and f). The fine-grained, light-grey to black mudstones commonly show evidence of strong soft-sediment deformation in the form of slump folds (Fig. 4g), micro-scale faulting, de-watering structures, ball-and-pillow structures, and load casts. The thickness of the mudstone layers ranges from a few mm up to several cm. The arkose and greywacke interbeds are pale in colour and composed of felsic volcanic material, dominated by quartz and plagioclase grains (Fig. 4f). In particular within the coarser beds, aphyric, quartz-phyric, feldspar-phyric, quartz-feldsparphyric rhyolite and dacite clasts can be observed. The sub-angular to sub-rounded grains vary typically between 0.02 mm and 1.6 mm in size. However, lithic fragments up to several cm in size are found, locally forming up to two-metre-thick conglomerate interbeds. Primary sedimentary structures such as load casts, cross bedding and graded bedding are well preserved and can be used for way-up determinations. 3.3. Conglomerate and breccias facies The most prominent type of conglomerate is characterised by felsic volcanic clasts (quartz-phyric, quartz-feldspar-phyric, feldspar-phyric and/or aphyric rhyolite to dacite), and less commonly mafic volcanic clasts (andesite to basalt), all derived from the nearby Skellefte Group volcanic rocks. Some volcanic clasts

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Fig. 4. Outcrop and micro photographs of the lithological units within the study area: (a) quartz-phyric rhyolite crosscut by polyphase quartz-calcite-vein, cross polarized light (XPol); (b) folded quartz-rich layer (S0 ) within calcite-quartz-mudstone, plane polarized light (PPol); (c) carbonate-cemented conglomerate with minor shale layers and distinct vertical stretching; (d) carbonate-cemented conglomerate, XPol; (e) alterning shale- and arkose to greywacke layers with rhyolitic boulder, soft-sediment deformation; (f) coarse grained, graded arkose (upper part) and fine grained argillite (lower part), foliation (S1 ) crosscuts bedding (S0 ), XPol; (g) Slump folding of alterning shale- and arkose-layers (turbidites); (h) rhyolitic to dacitic conglomerate, pale aphyric and quartz ± feldspar-phyric rhyolite clasts; (i) pebbly sandstone with angular shale

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Table 1 Lithologies in the Vargfors basin. Abbreviations: qz: quartz, (k-)fsp: (potassium-)feldspar, cal: calcite, plg: plagioclase. Facies

Characteristics

Carbonate-rich sedimentary rocks Fine grained, laminated cal1. Cal-qz-mudstone (>70% cal, <30% qz) and qz-rich (<30% cal, >70% qz) layers; strong deformation, qz-veining Clasts: pale grey; aphyric and 2. Polymict conglomerate with qz-, fsp, qz-fsp-phyric rhyolites carbonate-rich to dacites, altered rhyolites to dacites, qz- and fsp-fragments matrix Matrix: mainly recrystallised cal (80–95%), qz (20–5%); intense weathering Turbiditic unit Light to dark grey, laminated, 3. Mudstone soft sediment deformation (slump folds, micro scale faulting, ball-and-pillow structures, load casts), cross-bedding, layer thickness: few mm up to several cm Grey, grains: qz, plg, (±k-fsp), 4. Sandstone (arkose to greywacke) aphyric-, qz ± fsp-phyric rhyolite and dacite; normal grading, soft sediment deformation, cross bedding, layer thickness: 1 cm–2 m Conglomerate and breccias Clasts: pale grey; aphyric and 5. Brecciaconglomerate with qz-, fsp, qz-fsp-phyric rhyolites rhyolite–dacite lithic to dacites, (±dark grey andesties to basalts), qz- and clasts fsp-fragments Matrix: sandy to silty, qz- and fsp-grains; (±cal) Clast supported 6. Conglomerate Clasts: pale grey; aphyric and with rhyolite–dacite qz-, fsp, qz-fsp-phyric rhyolites lithic clasts to dacites, altered and/or mineralized rhyolites to dacites, qz- and fsp-fragments; (±dark grey andesties to basalts) Matrix: sandy to silty, qz- and fsp-grains; (±cal) Clasts: pale grey; aphyric and 7. Conglomerate with mudstone and qz-, fsp, qz-fsp-phyric rhyolites rhyolite–dacite lithic to dacites, altered rhyolites to clasts dacites, qz- and fsp-fragments; laminated turbidite fragments Matrix: sandy to silty, qz- and fsp-grains; (±cal) Clasts: pale grey; aphyric and 8. Polymict conglomerate with qz-, fsp, qz-fsp-phyric rhyolites rhyolite–dacite and to dacites, altered rhyolites to granite lithic clasts dacites, granodiorite to tonalite qz- and fsp-fragments; locally: dark grey andesites to basalts and reddish rhyolites to dacites Matrix: sandy to silty, qz- and fsp-grains; (±cal)

Grain size

Sorting

Rounding/crystal shape

Example

2–20 ␮m

Good

Rounded to sub-rounded

7220394, 1678845

1 mm–50 cm

Poor

Well-rounded to sub-rounded

7224396, 1675502

0.05–0.5 mm

Moderate

Subhedral to anhedral

2–20 ␮m

Good

Rounded to sub-rounded

7223384, 1676741

0.02–1.6 mm locally bigger

Good to moderate

Sub-rounded to sub-angular

7223670, 1676679

3 mm–15 cm

Poor

Sub-rounded to very angular

7224190, 1675993

2 ␮m–2 mm

Poor

2 mm–2 m bulk: 1–10 cm

Poor

Well-rounded to sub-angular

7224225, 1675993

2 ␮m–2 mm

Poor

2 mm–2 m bulk: 1–10 cm

Poor

Well-rounded to sub-angular Well-rounded to sub-angular

7223892, 1676791

2 ␮m–2 mm

Poor

2 mm–2 m bulk: 5–15 cm

Poor

Well-rounded to sub-angular Well-rounded to sub-angular

7221520, 1681000

2 ␮m–2 mm

Poor

are mineralized and/or sericite-altered. Sorting of the mainly clastsupported conglomerate is generally poor and clasts range from 0.5 mm to 2 m in size with a predominant size between 1 mm and 10 mm (Fig. 4h). The silty to sandy matrix consists mainly of well-rounded to sub-angular quartz and feldspar grains. Locally, this conglomerate is interbedded with sandstone, arkose and greywacke layers. In parts of the basin, angular clasts and boulders

Well-rounded to sub-angular

(3 mm–1 m) of the turbiditic unit constitute 5–95% of the clast population within the conglomerate (Fig. 4i and j). Locally, slump breccias consisting of sandstone-mudstone fragments with only a few rounded felsic clasts can be observed (Fig. 4k). Rhyolitic to dacitic breccia-conglomerates occur close to the Skellefte–Vargfors contact in compartments II and III, but also within the Skellefte Group volcanic sequence. This

clast; (j) conglomerate with volcanic and argillite clasts, PPol; (k) slump breccia with angular, laminated turbidite and round quartz-phyric rhyolite clasts; (l) rhyolitic to dacitic breccia-conglomerate, angular quartz ± feldspar-phyric rhyolite clasts; (m) tonalite clast from a conglomerate, partial pseudomorphic replacement of plagioclase by calcite, XPol; (n) polymict conglomerate, clasts: aphyric and quartz ± feldspar-phyric rhyolite, basalt, tonalite to granodiorite; (o) strongly quartz-phyric rhyolite.

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breccia-conglomerate is distinguished from the rhyolitic to dacitic conglomerate by the characteristic coexistence of rounded and very angular volcanic clasts (Fig. 4l). The N-part of compartment VI is dominated by a conglomerate with clasts derived from Skellefte Group rhyolites to dacites, and clasts of coarse-grained tonalite to granodiorite inferred to have been derived from the nearby Jörn intrusive complex (Fig. 4m). One further conglomerate type, previously described as the Dödmanberget conglomerate (Kautsky, 1957) additionally contains mafic and reddish felsic volcanic clasts, the latter originating from the Arvidsjaur Group (Fig. 4n). 3.4. Intrusive rocks Two types of undeformed sills and dykes locally intrude the sedimentary pile. The first type comprises dark, fine- and even grained ultramafic intrusions. The second type consists of pale, fine-grained felsic dykes with mainly rhyolitic to dacitic composition. 4. Stratigraphy within the Vargfors syncline 4.1. Compartment I Constrains to the compartment I stratigraphy are based on observations in the central part of the compartment. The lower part of stratigraphy is dominated by monomict conglomerates with and without mudstone lithic clasts and sandstones (Fig. 5). The uppermost part of the stratigraphy is dominated by a polymict granite-bearing conglomerate. Neither the contact between the monomict and polymict conglomerates nor the contact to the underlying Skellefte Group volcanic rocks is exposed. 4.2. Compartment II Undisturbed successions in the SE-part of compartment II show that carbonate-cemented conglomerates occur mainly in the lowermost parts of the stratigraphy (Fig. 5). They rest unconformably above Skellefte Group volcanic rocks. The carbonate-cemented conglomerates grade into the monomict conglomerates with rhyolite–dacite lithic clasts. Intercalations with turbiditic rocks and sandstones are common and show gradational contacts towards the monomict conglomerates. The monomict conglomerate is overlain by rocks of the turbiditic unit, whereas the contact is not exposed. Stratigraphy in the northernmost part of the compartment is not known. 4.3. Compartment III The stratigraphical succession within compartment III shows similarities to compartment II with lower parts of stratigraphy dominated by monomict conglomerates with intercalations of carbonate-cemented conglomerates (Fig. 5). Contrasting to compartment II, the lowermost part of compartment III is composed of a breccia-conglomerate with rhyolite–dacite lithic clasts resting unconformably on top of the Skellefte Group volcanic rocks. The uppermost exposed parts of stratigraphy within compartment III comprises turbiditic rocks with abundant conglomerate interbeds resting conformable on top of the monomict conglomerate. 4.4. Compartment IV The lowest part of stratigraphy in compartment IV is exposed along its faulted N-contact and characterised by highly strained carbonate-cemented conglomerates and monomict conglomerates (VI-N in Fig. 5). They are overlain by rocks of the turbiditic unit. The undisturbed succession in the central and S-parts of compartment IV shows monomict conglomerates grading into turbiditic

173

rocks with frequent conglomerate and sandstone interbeds. The turbiditic unit is grading upwards in stratigraphy into monomict conglomerates with minor turbiditic beds. The S-contact of compartment IV towards the Skellefte Group volcanic rocks is not exposed. 4.5. Compartment V Stratigraphy within compartment V is dominated by abundant occurrence of turbiditic rocks with minor intercalations of rhyolitic–dacitic conglomerates with mudstone clasts. A 50 cm thick bed of carbonate-cemented conglomerate was observed on outcrops along a high-strain zone at the S-contact of the compartments, as well as in a bore hole intersecting the contact. The N-contact of the compartment is not exposed. 4.6. Compartment VI The S-part of compartment VI is characterised by turbiditic rocks grading into monomict conglomerates (VI-S in Fig. 5). Intercalations of sandstones and mudstone-bearing conglomerates are common within the monomict conglomerate. The contact towards the Skellefte Group volcanic rocks to the S is faulted and associated by calcite-rich mudstones and conglomerates. The N-part of compartment VI is dominated by a polymict conglomerate resting unconformably above aphyric Skellefte Group rhyolites (VI-N in Fig. 5). 5. Structures within the Vargfors syncline The central part of the Vargfors basin occurs within an ESE–WNW-trending regional-scale syncline and has an average width of approximately 2 km. Two sets of interlinked high-strain zones with approximately SE–NW and approximately NE–SW orientations, define specific fault-bound compartments (Fig. 3). The main synclines within the fault-bound compartments are upright to overturned, variably asymmetric and their hinges are sub-parallel to the SE–NW-striking syncline-bounding high-strain zones. A regional axial culmination occurs along the compartment II/III contact: the synclines plunge gently to the NW and SE, to the W and E of this contact, respectively. The observed fold axes are typically sub-parallel to the basin-scale syncline hinges, but locally show highly curvilinear shapes. This is attributed to sub-vertical stretching during the crustal shortening, as demonstrated by the sub-vertical mineral lineations occurring both along the main foliation surfaces, and as down-dip lineations along the SE–NW faults (Fig. 6a and b). The steeply dipping, approximately SSE–NNW-striking main foliation transects both limbs of the basin-scale synclines with the same vergence (Fig. 3), although it is locally aligned parallel to the axial surfaces of the meso-scale folds (Fig. 6d). The foliation intensity is typically lowest in the cores of the synclines, where it cuts flat-lying bedding surfaces at a high angle (Fig. 6c). Within the higher-strain domains, axial surfaces of tight to isoclinal folds (Fig. 6e), and the related foliations, generally strike parallel to the main high-strain zones, thus deviating from the regional ∼SSE–NNW trend. This is especially the case for the major SE–NW syncline-bounding high-strain zones, but in lesser extend also for the NE–SW-striking high-strain zones. 5.1. Compartments I and II The few observed bedding planes in compartment I dip gently towards NW and show evidence for open folding. Compartment II is characterised by strata with N-W dips, variably modified by open, upright folds with SE–NW-striking axial surfaces (Fig. 7a).

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Fig. 5. Stratigraphic columns of fault-bound compartments in the central parts of the Vargfors syncline. The roman numerals (I–VI) refer to specific fault-bound compartments (c.f. Figs. 2 and 3). Compartments IV and VI were divided in northern (IV-N and VI-N) and southern (IV-S and VI-S) stratigraphic columns to emphasise the differing contact relationships.

The approximately SE–NW-striking main foliation is locally transposed into approximately WSW–ENE orientation, particularly in the vicinity of the WSW–ENE-striking fault in the NW part of the compartment, where high localised strain is further reflected by overturned beds along the compartment II/III contact. The approximately WSW–ENE-striking high-strain zone is associated with an upright anticline cored by a strongly quartz-phyric volcanic rock, which conformably underlies the Vargfors Group monomict conglomerates and breccia-conglomerates (Fig. 7a). The anticline marks the regional axial culmination, therefore implying this location shows the deepest exposure through the syncline in the area.

5.2. Compartment III The sedimentary successions within this highly strained compartment are intensely folded and faulted. The strata show tectonic thickening, repetition and overturning towards SW, particularly within the central part of the compartment, which shows highest strain (Fig. 7b). More locally, the greatest strains were apparently localised into the rheologically weak carbonate-cemented conglomerate facies, as recorded by their intense sub-vertical mineral lineations, and transposition of adjacent rocks into sub-vertical attitudes (Figs. 6g and 7b). Sub-vertical tectonic transport is also reflected by the distribution of the measured fold axes (Fig. 7b), which scatter along a sub-vertical, SE–NW-trending girdle on the stereographic projection (Fig. 7b). Local refolding of the overturned units indicates progressive deformation within the high-strain domain (Fig. 6f–i). Lower-strain domains occur both locally within the high-strain domains (Fig. 6g), and contrary to most of the other compartments, close to the compartment margins. The main foliation cuts the structural grain of the compartment at an angle of approximately 15◦ (Figs. 3 and 7b), whereas the volcanic rocks bounding the compartment, both in the N and S, show only weak foliation. At the compartment II/III contact, the main foliation follows the approximately NNE–SSW orientation of the high-strain zones and the related overturned folds, and overprints the WSW–ENE-trending anticline that is parallel to the NNE–SSWstriking high-strain zone.

5.3. Compartment IV Structures within compartment IV indicate increasing strain from an open syncline in the SW to faulted and overturned structures in the NE (Fig. 7c). High strain along the NE-margin of the compartment is manifested by shear fabrics in the carbonatecemented conglomerates along the contact, and overturning of the sedimentary strata within 250 m of the contact. A progressive decrease in fabric intensity within the bounding plagioclase-phyric volcanic rocks from strongly foliated along the contact, to unfoliated approximately 10 m away from it, further implies that the highest strain was partitioned into the sedimentary rocks and the sediment-volcanic contact. Thus, the NE-margin of the compartment is interpreted as a set of parallel, NE-dipping high-strain zones (Fig. 7c). Localised overturning occurred also along discrete reverse high-strain zones within the N-half of the compartment (Fig. 6h). Deformation localisation took place along the approximately NE–SW-striking high-strain zones, as shown by folding adjacent to the high-strain zones, deflection of structures into the high-strain zones, and strong foliations occurring parallel to the high-strain zones. These foliations within the basin-bounding volcanic rocks are more intense than their reverse high-strain-related equivalents, and continue to the NE, beyond the sediment-volcanic contact. An open, upright, and gently SE-plunging syncline characterises the S-half of the compartment. In the core of the syncline, high-strain zones occur preferentially within carbonate-cemented conglomerates. However, these high-strain zones are not continuous, and therefore differ from those along the NE-margin of the compartment. Strain decreases towards S and SE, where the structure is defined by gently SE-dipping, virtually undeformed stratified sedimentary rocks. 5.4. Compartment V Bedding planes in compartment V are folded into an open, SE-plunging syncline, with domains of nearly undeformed beds dipping gently SE in the core of the syncline. Transition from the SE dips into the typical SE–NW-strike is seen along localised, high-strain-related folds, as illustrated in Fig. 6i. Closer to the

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Fig. 6. Outcrop photographs and sketches of the characteristic deformation structures within the study area: (a) Sub-vertical mineral lineation in sandstone. Vertical section. (b) Stretching lineation defined by re-crystallized quartz grains. S-margin of compartment V. Section dips steeply towards south. (c) Sub-horizontal bedding surfaces perpendicularly cut by tectonic main foliation. Vertical section. (d) Open, upright, gently SE-plunging F1 folds deform layered sedimentary rocks. (e) High-strain folding close to the NE-margin of compartment IV. (f) Detailed profile through an outcrop with refolded, overturned layered rocks. (g) A detailed profile through an outcrop illustrating the abrupt change from right-way-up strata in the hinge of an open fold into overturned layering against its southern limb. Symbols as in (f). (h) Outcrop-scale overturned

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Fig. 7. Geological profiles through compartments II–VI, and equal-area, lower-hemisphere stereographic projections of the related structural measurements (S0 = bedding, S1 = main foliation, F1 = fold axis, L = lineation). (a) Compartment II and III-W, (b) compartment III-E with close-up, (c) compartment IV, (d) compartment V, (e) compartment VI, (f) longsection W and E, (g) all measured lineations. See Fig. 2 for legend.

compartment IV/V contact, these high-strain zones cause stronger transposition, nonetheless never penetrating through the whole compartment. Pronounced shearing was recorded along the SWcontact of the compartment (Fig. 7d). The character of this

high-strain zone changes considerably along strike; in the SE, the strata were transposed and overturned on the NE-side of the thin, distinct high-strain zone, which contains tectonic inclusions of rhyolite and limestone (Fig. 6j). Further to the NW, the turbiditic

folds with axial surface-parallel main foliation. Vertical section. (i) Detailed profile illustrating the development of a fault-related anticline occurring in the transition from ∼undeformed, gently SE-dipping strata into D1 -transposed strata striking SE–NW. (j) Outcrop sketch and related profile showing the contact relationships between the sedimentary and volcanic rocks along the SW-boundary of compartment V. Note the increasing intensity of main foliation towards the shear zone. I = A remnant sliver of rhyolite, and II = fine-grained, intensely foliated marble within the shear zone.

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Fig. 7. (Continued)

sequences dip gently, although an intense foliation was developed parallel to the high-strain zone in the bounding volcanic rocks, similar to those in Fig. 6j. In drill core, this high-strain zone could be identified as a 50–100 cm wide zone of highly deformed carbonatecemented conglomerate. A small massive sulphide deposit (Åliden) is situated in a high strain zone, close to the contact between the Skellefte and Vargfors Groups in the N-part of the compartment. In the vicinity of the Åliden fault (Figs. 2 and 3), Skellefte Group volcanic rocks and Vargfors Group sedimentary rocks show hydrothermal alteration associated to the mineralization. 5.5. Compartment VI The pattern defined by the structural form lines on the geological map and cross-section (Figs. 3 and 7e) indicate strain partitioning into a major high-strain zone along the SW-margin of the compartment, where NE-side-up kinematics were observed. On the contrary, the sedimentary units farther to the NE experienced folding under low-strain. The sedimentary-volcanic contact along the NE-margin of the compartment is a stratigraphic unconformity. No axes of meso-scale folds were measured, but the calculated ␤-axis plunges gently towards SE (Fig. 7e), similar to the compartments IV and V. The main foliation is sub-parallel to the syncline-bounding high-strain zone in the SW, and to the NE shows a NNW–SSE strike oblique both to the margins and the axis of the syncline (Fig. 3). 6. 2D forward modelling 6.1. Modelling methods Two-dimensional forward modelling along vertical profiles through selected compartments was performed in order to test

how the described geometries and lithostratigraphies can be generated. The 2D-MOVETM software package (Midland Valley) was used to test if the observed geometries can be obtained by the suggested extensional and compressional evolution (Allen et al., 1996) in balanced cross-sections (c.f. Dahlstrom, 1969; Gibbs, 1984, 1990). Modelling was performed in the compartment VI due to its simple geometry, its low strain and its well exposed boundaries. Furthermore, the results from compartment VI were utilised for modelling of compartment IV, which is also well exposed with a well constrained structural geometry but shows higher strain than compartment VI. The results of 2D-modelling aid the estimation of the required amount of extension, compression and primary fault geometries.

6.1.1. Compartment VI During the extensional phase of the basin evolution, different geometries for the listric normal faults were tested in a model comprising a package of undisturbed horizontal layers down to ∼2200 m depth, representing Skellefte Group volcanics rocks. The depth-extent was varied between 1000 m and 2500 m, the dip angle between 60◦ and 90◦ at surface, and the curvature (c.f. Gibbs, 1984, 1990). The “simple shear” and “fault parallel flow” algorithms were tested in combination with varying shear angles, fault geometries and sedimentation rates in order to simulate sedimentation and sub-basin formation along the normal growth faults (Fig. 8a). During the transitional phase between crustal extension and compression, erosion was tested on several elevation levels ranging from −50 m to −300 m. Several layers, representing the granitebearing polymict conglomerate, were added manually on top of the turbiditic sedimentary rocks after the erosional event (Fig. 8b). This resulted in an angular unconformity between the underlying

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Fig. 8. 2D-forward model for compartments VI and IV. Compartment VI: (a) end of the extensional phase, (b) erosion and deposition of the polymict conglomerate, (c) compressional phase, fault inversion and transposition of strata towards the fault, (d) folding. Compartment IV: (e) end of the extensional stage, (f) end of fault inversion and folding, (g) formation of break-back fault and juxtaposition of volcanic rocks, (h) ongoing formation of break-back faults, (i) rotation of the inverted normal fault, (j) formation of break-back faults.

turbiditic rocks and monomict conglomerates and the overlying polymict conglomerates. Crustal shortening was simulated by compressing the modelled sub-basin in increments of 50 m of slip along the faults. During the modelling, strain was accommodated by the listric fault, and both the simple shear algorithm and the fault parallel flow algorithm with varying shear angles were applied. A fault inversion phase (Fig. 8c) was followed by a folding phase (Fig. 8d), which was simulated using the “unfold to target” tool in MOVE (using a symmetric buckle-fold-shaped template) in order to perform folding manually. Use of the described method in forward modelling is justified by analogue modelling supporting initiation of basin inversion via fault reactivation when (1) a rheologically weak lithological unit, such as carbonate-rich sediments occur at the base of the sub-basin (Del Ventisette et al., 2006), or (2) continued sedimentation on the normal fault leads to flatter dip of the fault at

depth, hence enhancing the fault reactivation (Pinto et al., 2010). Furthermore, the observed foliations along the NW–SE-striking high-strain zones and partly along the NE–SW-striking high-strain zones reflect localised deformation rather than penetrative and distributed shortening affecting both the basement and the basin-fill succession. Down-dip lineations along the inferred reactivated normal faults and the apparent lack of coeval strike-slip deformation infer that the bulk shortening at the onset of basin inversion was orthogonal to the syncline margins, hence allowing us to orient the profiles used in forward modelling at right angles across the syncline axial traces, i.e. parallel with geological profiles (Figs. 2 and 3). 6.1.2. Compartment IV Modelling of the extensional phase in compartment IV was performed according to that in compartment VI (Fig. 8e). As the polymict conglomerate and the associated unconformity are not exposed in compartment IV, also the erosional phase was neglected.

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The compressional phase commenced with a phase of fault inversion, followed by folding (Fig. 8f), comparable to the inversion of compartment VI. Tight folding and overturning of the sedimentary strata close to the compartment margin point towards significant reverse movements along break-back faults (Figs. 6h and 8g–i) post-dating the initial normal fault inversion. Alternatively, the normal fault may have been subjected to pronounced rotation, and finally overturning of the normal fault (Fig. 8i). Irrespective of the mechanism, the cumulative dip-slip displacement between the Skellefte Group volcanic rocks and the basin fill along the NE compartment margin is significant. Therefore, two alternatives have been tested for the N-margin of the compartment. 6.2. Results and evaluation The results best corresponding to the field observations were reached by using a steep listric fault with a detachment depth at ∼2000 m (Fig. 8a and e). Extension was performed in 50 m steps of slip along the fault, utilising the simple shear algorithm and a positive shear angle of 75◦ , until a cumulative movement of 1500 m was reached. Sediments were added automatically at −20 m elevation to represent an almost complete basin fill. The result shows sedimentary sub-basins with ∼4600 m width and ∼1600 m depth (Fig. 8a and e), corresponding to approximately 18% extension. In compartment VI, two layers with 100 m thickness each, representing the granite-bearing polymict conglomerate, were added manually on top of an erosional surface at −250 m elevation. In the compressional phase, the simple shear algorithm and a positive shear angle of 85◦ gave the best results for fault inversion. Modelling showed that the listric fault has to be steep already early during inversion in order to create a strong transposition of strata towards the fault and the distinct asymmetry of the syncline (Fig. 8c and f). The use of moderate and gently dipping faults resulted in a less intense transposition of strata towards the fault and was therefore not satisfactory. After 650 m of cumulative movement along the fault, corresponding to approximately 9% of shortening, folding was simulated manually. The resulting sub-basin in compartment VI has an approximate width of 2600 m and a depth of 400–500 m (Fig. 8d). Modelling in compartment IV showed that both alternatives, either the formation of break-back faults alone (Fig. 8g and h), or rotation of the inverted normal fault later followed by formation of break-back faults (Fig. 8i and j), may have resulted in geometries that may be correlated with the field observations. Moreover, both alternatives require that the break-back faults need to have a steep attitude in order to create a sub-vertical transposition of strata towards the faults (c.f. Fig. 6h). 7. Model for the restoration of the central part of the Vargfors syncline 7.1. Tentative 2D-model for the restoration of the remaining compartments Based on the results of two-dimensional forward modelling we propose a model for a schematic restoration of the remaining faultbound compartments in the central parts of the Vargfors syncline. 7.1.1. Compartments II and III Open folding in compartment II indicates that this part of the syncline has undergone only weak shortening. Tectonic movements along the NE–SW-striking high-strain zones in the W-part of compartment III (Fig. 9a and b) caused doming of the Skellefte Group volcanic rocks prior to the observed overturned folding. Reactivation of the early normal fault, followed by fault-splay

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formation, locally created overturned folding. However, the overturned strata occur away from the SW-margin of the compartment, which implies that the high-strain zones in profile III-E are continuous across compartment III, but show a clockwise-rotation towards E in the map view. Deformation in the central and the E-parts of compartment II is dominated by fault inversion along the N-contact of the syncline. The presence of overturned bedding in large parts of compartment III would suggest that there originally were several normal faults (Fig. 9c). It is likely that the oldest faults became flatter due to progressive extension, and were therefore more prone to reactivate during the basin inversion compared to younger and steeper faults (Bailey et al., 2002). The largely overturned strata within the southernmost subblocks (Fig. 9c) was probably produced by (i) fault reactivation, and (ii) subsequent rotation and overturning of the faults during progressive shortening, and (iii) development of normal-fault-parallel break-back faults during the last stages of compression. The absence of carbonate-rich sediments along the NW-margin of compartment III is either due to (1) the absence of it in the NW-part of the syncline, reflecting the complex evolution of the extensional sub-basin with the formation of several normal faults, which influence sedimentary development, or (2) a break-back fault that displaced the marginal carbonate-rich conglomerate, and juxtaposed the quartz-phyric felsic volcanic rocks against mudstones and sandstones. 7.1.2. Compartment V The division of the syncline into two sub-synclines with a basement high in-between (Fig. 9d) could be obtained by two N-dipping major normal faults with an antithetic S-dipping fault in between. The greatest strain was localised into the NW-part of the syncline. Only the lower most part of the stratigraphy was preserved due to NE-side up reverse faulting, similar to that along the N-margin of compartment IV. The present geometry of the NW sub-basin is consistent with a NW-plunging local syncline, bound by the reactivated normal fault and a footwall shortcut-fault (Fig. 9d). The Åliden massive sulphide mineralization, which formed at the contact between the Skellefte and Vargfors Groups, was transposed into the reactivated normal fault. 7.2. 3D-model of the central part of the Vargfors basin 2D-modelling of compartments VI and IV showed that the results are compatible with the previously suggested structural succession comprising of initial extension followed by later inversion (Allen et al., 1996). The different and opposing fault geometries and the corresponding fold asymmetries between the compartments may therefore be attributed to variable normal fault geometries due to location on the opposing limbs of the present-day synclinal structure (Fig. 10). Consequently, the approximately NE–SW-striking high strain zones must have formed synchronously with the extensional normal faults, i.e. as transfer faults linking together the normal faults (c.f. Gibbs, 1984). Recognition of this pattern with two synchronously active sets of faults allows us to suggest a tentative 3D-model explaining the evolution of the central parts of the Vargfors syncline/basin. Evolution of the approximately WNW–ESE basin-bounding faults could be deduced from the syncline asymmetry, degree of tectonic transposition associated with the faults, and kinematic observations in the field and from the thin sections. These constraints were further evaluated by 2D-modelling. In contrast, little is directly known about the characteristics of the transfer faults and, for this reason, their kinematic evolution needs to be inferred indirectly. Firstly, the oblique trend of the overprinting regional foliation may reflect transpressional deformation with

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Fig. 9. Tentative 2D-model for the restoration of compartments II, III and V showing the geometry during extension, start of compression (except II and III-W) and the present geometry. (a) Compartment II, (b) compartment III-W, (c) compartment III-E, (d) compartment V. Note the varying scale between extension/start of compression and present geometry illustrations.

approximately SW–NE crustal shortening during the basin inversion, which would be resolved as dextral strike-slip movement along the transfer faults. Secondly, since the transfer faults were developed synchronously with the normal faults (this paper) they must have accommodated the variable strain and direction of tectonic transport between the compartments. For this reason, a generally valid movement sense for the transfer faults may not be defined. A careful inspection of the spatial distribution of rock types on the geological map (Figs. 2 and 3) might give information about the sense of movement along the transfer faults. However, since the present-day configuration is a combined result of crustal extension and the overprinting crustal shortening, the effect of these tectonic events may not be isolated from each another. The best possibility to investigate the nature of the transfer faults is to look at the sedimentary strata in their vicinity, in lowstrain domains in the syncline cores where the inversion of the normal faults had the least effects. Gently SE-dipping bedding planes in such low-strain domains in compartments III, IV, V, and VI, and NW-dipping bedding planes in compartments I and II, is attributed to their pre-folding attitudes, thus inferred to have generated during the crustal extension. This pre-compressional dip of bedding planes is inferred to result from tilting of the faultbound blocks due to syn-extensional tectonic displacements along the transfer faults during the progressive opening of the sub-basin towards NW and SE (Fig. 10). The opening of the sub-basin resulted in an extensional displacement gradient along the main structural grain, where the compartments within the early-formed parts of the sub-basin (compartment II/III boundary) became deeper than the later-formed ones. Consequently, the dip-slip component during the crustal extension is attributed to the progressive opening of the basin, whereas the strike-slip component is more unsystematic and of a more local nature, resulting from accommodating the variations in deformation between the individual neighbouring compartments. However, intense foliations along the contact between compartments II–III and IV–V show that at least the largest of the transfer faults were affected by compressional deformation.

The syn-extensional geometry of the sub-basin is attributed in part to lithological and structural inhomogeneities within the Skellefte Group volcanic rocks that act as a basement for the Vargfors Group sedimentary rocks, and in part to the progressive rifting, starting at the compartment II/III contact and producing different geometries within the different compartments. Slump brecciation and intense slump folding confirm the rapid subsidence during ongoing sedimentation in the early-formed compartments. Pronounced soft-sediment deformation indicates that highest displacements took place along the normal growth faults. Subsequent inversion was inferred to lead to a similar displacement gradient but with an opposite movement sense, eventually resulting in stronger uplift of the central part compared to the distal parts of the sub-basin. This may have caused additional tilting of the central blocks and contribute to the formation of the regional axial culmination. The basin-scale curvilinear nature of the syncline axis (F2 in literature; e.g. Bergman Weihed, 2001) is attributed to the initial dip of strata away from the core of the syncline, combined with subsequent tilting of the fault blocks, rather than due to Nstriking steep, mainly dip-slip cross faults and N-striking open to tight folds as suggested by Allen et al. (1996). The inversion of sedimentary basins results predominantly in the formation of antiforms rather than synforms as seen in experiments and field observations (Bailey et al., 2002; Panien et al., 2005). Therefore, the formation of synforms during inversion of the Vargfors basin probably required certain specific conditions, such as gliding along (i) basal, carbonate-rich mudstones and carbonatecemented conglomerates (comparable to the ductile layer in Del Ventisette et al., 2006) and/or (ii) a zone with high fluid pressures related to hydrothermal fluid circulation (De Wit, 1982). The carbonate-rich mudstone and the carbonate-cemented conglomerates observed in the study area originate either from the deposition of carbonate-rich sediments during an early stage of sedimentation or hydrothermal alteration of conglomerates and mudstones, seen as pseudomorphic replacement of plagioclase by calcite especially in the matrix of the altered conglomerates (Fig. 4d), but also in certain layers of the mudstones (Fig. 4f). Such hydrothermal alteration

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Fig. 10. Simplified 3D block model of compartments II–VI, showing the post-extensional geometry of the central parts of the Vargfors sub-basin. Note the opposing locations of the basin-bounding faults within the specific compartments and their primary dips away from compartment II/III contact. Sedimentary rocks are excluded for clarity.

could be related to the major VMS forming event in the uppermost part of the Skellefte Group (Allen et al., 1996). Nevertheless, more strain was partitioned into the carbonate-rich lithologies, which now occur as deformed, carbonate-rich lenses of rock along the inverted compartment contacts. Break-back and short-cut faults (Figs. 8 and 9c and d) that formed during progressive shortening, resulted in fault-induced folding in the vicinity of the faults, as well as the formation of low-strain domains between the faults. No regionally valid fold-vergence direction could be determined, inferring that the overturning of the folds was controlled by the local fault geometry instead. 8. Discussion Rocks within the Vargfors syncline show low strains compared to most areas in the Skellefte district, which makes it possible to infer the three-dimensional pattern of reactivated normal faults, transfer faults and later formed break-back and short-cut faults. Despite the higher strain, a strikingly similar fault pattern occurs in

other parts of the Skellefte district, and is inferred to be of the same origin. For example, the Deppis-Näsliden shear zone and VidselRöjnoret shear system might be early extensional faults similar to transfer faults in the Vargfors syncline, as promoted by Skyttä et al. (2010). Carranza and Sadeghi (2010) and Skyttä et al. (2010) suggested that the Skellefte district formed as a straight system, implying that the Deppis-Näsliden shear zone and the VidselRöjnoret shear system could have acted as rotation axes causing rotation of the Kristineberg and Boliden-blocks with respect to the central parts of the Skellefte district. The main compressional deformation created the upright folds with NW–SE-trending axial surfaces, as well as down-dip lineations on the reverse faults, which might indicate SW–NE-compression perpendicular to the normal faults. This deformation pattern is consistent with the interpretation by Weihed et al. (2002a) who attributed SW–NE-compression to accretion of the Skellefte volcanic arc to the Archean craton during the Svecofennian orogeny. Later rotation of the bulk compression into WSW–ENE-orientation resulted in the oblique foliation pattern in the Vargfors syncline, where foliation crosscuts both fold limbs, oblique to the main

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structural grain of the sub-basin. Alternatively, sinistral transpressional conditions could have caused fault reactivation and the development of the oblique foliation pattern in the Vargfors syncline. Further studies are required to unravel the bigger scale tectonic framework of the region. Subsequent rotation of the stress field terminates in the E-W-compressional event, reported by Bergman Weihed (2001), which is the result of the oblique collision of Fennoscandia with Sarmatia during the Svecobaltic orogeny after the accretion of the arc system to the Archean craton (Weihed et al., 2002a, 2005; Lahtinen et al., 2004, 2005). Even though no structures related to that later E-W shortening (D3 in Bergman Weihed, 2001) are obvious in the study area, evidence can be observed outside the study area, seen as crenulation cleavage in mainly mudstones (Dumas, 1986; Bergman Weihed, 2001) and as movement along the Deppis-Näsliden shear zone and Vidsel-Röjnoret shear system. Despite this crenulation cleavage, no related folding can be observed in the area, which suggests that a significant proportion of this later deformation was localised into major shear zones. Bergman Weihed (2001) proposed that the Deppis-Näsliden shear zone and the Vidsel-Röjnoret shear system formed under this later E-W compressional deformation. However, considering the results of the present study, it may be strongly inferred that these shear zones have their origin in the crustal extension, and were only reactivated during this later deformation (c.f. Skyttä et al., 2010). VMS mineralization in the Skellefte district occurred during extension of an immature continental volcanic arc (Allen et al., 1996, 2002; Weihed et al., 2005), with the early normal faults and, at least in parts, the transfer faults behaving as conduits for magmas and hydrothermal solutions. This can be observed for example as extensive hydrothermal alteration along the Åliden fault (Fig. 2), being most intense at the intersection of the normal fault with the Åliden-transfer-fault. Later inversion caused transposition of the massive sulphide bodies into steep orientation along reactivated faults (e.g. Åliden deposit in Fig. 9d) and/or into thrust and folded lenses as along the lower structural levels in the Kristineberg deposit (Årebäck et al., 2005). This fault activity implies fluid flow along the reactivated faults, which can result in a certain degree of remobilisation of ore minerals as observed for example in the Långdal deposit (Weihed et al., 2002b). 9. Conclusion The three-dimensional geometry of the sedimentary Vargfors syncline is the result of Palaeoproterozoic extension and subsequent compression. Early extensional NW–SE-striking normal faults and interlinked NE–SW-striking transfer faults define distinct fault-bound compartments, each with an individual structural geometry and stratigraphy. Constraints on the deformation style and mechanisms achieved by 2D-forward modelling are in agreement with the previously inferred inversion of the early normal faults during a regional crustal shortening event, associated with folding of the sedimentary rocks. A rheologically weak carbonaterich layer at the base of the sedimentary sequence favoured the fault inversion over more distributed shortening as the controlling deformation mechanism. Continued inversion led to progressive steepening of the faults, subsequent formation of break-back faults and associated local overturning of sedimentary strata. Acknowledgements This work is part of the VINNOVA 4D-modelling project and the PROMINE project, financed by VINNOVA and Boliden Mineral

AB and the European Union, respectively. Dr. Dave Coller and Dr. Peter Sorjonen-Ward are acknowledged for very constructive discussions and ideas. Constructive comments from journal reviewers Michael Stephens and Rick Squire are much appreciated and lead to a substantially stronger manuscript. Midland Valley Exploration Ltd. is acknowledged for providing the MOVETM software package under the “Academic Software Initiative”, as well as for support from Midland Valley’s geologists Dr. John Grocott and MSc, Jenny Ellis. The Geological Survey of Sweden and Geovista AB are thanked for contributions.

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