Multiple reactivation and strain localization along a Proterozoic orogen-scale deformation zone: The Kongsberg-Telemark boundary in southern Norway revisited

Precambrian Research 265 (2015) 78–103

Contents lists available at ScienceDirect

Precambrian Research journal homepage: www.elsevier.com/locate/precamres

Multiple reactivation and strain localization along a Proterozoic orogen-scale deformation zone: The Kongsberg-Telemark boundary in southern Norway revisited Thomas Scheiber a,∗ , Giulio Viola a,b , Bernard Bingen a , Max Peters c , Arne Solli a a

Geological Survey of Norway, 7491 Trondheim, Norway Department of Geology and Mineral Resources Engineering, Norwegian University of Science and Technology, 7491 Trondheim, Norway c Institute of Geological Sciences, University of Bern, 3012 Bern, Switzerland b

a r t i c l e

i n f o

Article history: Received 1 October 2014 Received in revised form 27 January 2015 Accepted 10 March 2015 Available online 23 March 2015 Keywords: Precambrian deformation zone Sveconorwegian orogen Kongsberg lithotectonic unit Structural analysis Zircon U–Pb geochronology Multiple reactivation

a b s t r a c t Structural analysis defines a multiphase Sveconorwegian tectonic evolution for the boundary zone between the Kongsberg and Telemark lithotectonic units in S Norway, referred to as the KongsbergTelemark Boundary Zone (KTBZ). This large-scale weakness zone developed predominantly within and at the margin of a c. 110 km long granitic belt, the intrusion of which is dated between 1170 ± 11 and 1146 ± 5 Ma by U–Pb SIMS zircon geochronology. The oldest KTBZ ductile fabric formed during the Sveconorwegian orogenic cycle (c. 1140–900 Ma) as a penetrative top-to-the-W shear fabric, which was subsequently reactivated selectively by sinistral transpression that formed characteristic mylonitic shear zones within the granitic belt. Later folding affected the area at the northern end of the Kongsberg lithotectonic unit. Analysis of the subregional foliation trajectories unravels the occurrence of a large-scale fold structure, the “Norefjell-Hønefoss Fold”. All these structures are in turn cut by late-Sveconorwegian, E-dipping shear zones and normal faults, which accommodated a distinct phase of exhumation of the Telemark lithotectonic unit in the footwall of the KTBZ. This extensional detachment widens toward the north, where it might have controlled the emplacement of the late-orogenic Flå granite. Since late Sveconorwegian times, the KTBZ was repeatedly reactivated in a brittle fashion forming complex fault patterns, extensive quartz vein networks and leading to the generation of the so-called “Store Rivningsbreksje”, a 100 km long brittle fault zone that follows the trend of the KTBZ and that locally juxtaposes blocks with different ductile precursor histories. The newly established deformation history helps to refine existing models for the orogenic evolution of the central Sveconorwegian orogen. The characterization of the Norefjell-Hønefoss fold structure provides a new perspective on Sveconorwegian geometries and fabrics in the area. The reactivation history established for the KTBZ helps to better understand the dynamics of long-lived weakness zones of Precambrian origin in general. © 2015 Elsevier B.V. All rights reserved.

1. Introduction Deformation in the Earth’s crust is commonly localized within relatively few, narrow and generally long-lived accommodation zones such as ductile shear zones or brittle faults (e.g., Sibson, 1977, 1986; Scholz, 1988; Govers and Wortel, 1995; Platt and Behr, 2011a). Repeated strain localization in these zones is facilitated by mechanical weakening (Prucha, 1992; Rutter et al., 2001), whereby any given volume of already deformed rock will take up the strain increments of later deformational episodes by reactivation of

∗ Corresponding author. Tel.: +47 7390 4461. E-mail address: [email protected] (T. Scheiber). http://dx.doi.org/10.1016/j.precamres.2015.03.009 0301-9268/© 2015 Elsevier B.V. All rights reserved.

pre-existing deformation zones, rather than by formation of new ones (e.g., Holdsworth et al., 1997, 2001; Viola et al., 2012). Reactivation may be accommodated by: (1) Ductile reactivation of early low-grade brittle faults, which generally become fully obliterated during the process (e.g., Passchier et al., 1990; Mancktelow and Pennacchioni, 2005), (2) ductile reactivation of earlier high-grade ductile fabrics due to a change in the governing stress field (e.g., Simpson et al., 2001; Occhipinti and Reddy, 2004), (3) brittle reactivation of ductile shear zones (e.g., Ziegler, 1996; Bezerra et al., 2014) and/or (4) brittle reactivation of earlier brittle features (e.g., Mirabella et al., 2004; Madritsch et al., 2009). Being able to unravel in detail deformation histories reflecting any possible combination of the reactivation modes above is of great importance particularly for the successful analysis of ancient orogens, where the preserved geological record is the summation of multiple geological events.

T. Scheiber et al. / Precambrian Research 265 (2015) 78–103

79

Fig. 1. (a) Sketch tectonic map of Scandinavia highlighting the areas affected by the Sveconorwegian orogeny. The northern boundary of the Sveconorwegian imprint is drawn following Røhr et al. (2013). (b) Tectonic framework of the Sveconorwegian orogen (modified after Viola and Henderson, 2010) and of the study area (framed in black) showing the main Sveconorwegian lithotectonic units, major Sveconorwegian ductile shear zones, Sveconorwegian late- to post-orogenic (c. 990–910 Ma) plutons, the “Store Rivningsbreksje” (“Great Friction Breccia”) according to Bugge (1928) and the zone identified as the Kongsberg-Telemark Boundary Zone (KTBZ) by the present study. Rectangle indicates location of Fig. 2. Abbreviations: SFDZ, Sveconorwegian Frontal Deformation Zone; s.z., shear zone.

The Proterozoic Sveconorwegian orogen in SW Scandinavia hosts several orogen-scale shear zones that record evidence for multiple reactivations, such as, for example, the Mandal-Ustaoset line (Sigmond, 1985), the Mylonite Zone (e.g., Stephens et al., 1996; Viola et al., 2011) or the Kristiansand-Porsgrunn shear zone (Mulch et al., 2005). These shear zones bound lithotectonic units or crustal blocks (Fig. 1; Andersen, 2005) with differing Sveconorwegian orogenic histories (see Bingen et al., 2008b and Roberts and Slagstad, 2014 for reviews and references). They record multiple Sveconorwegian deformation resulting from the collision between the SW Fennoscandia margin and another major plate, supposedly the continental plate Amazonia (Romer, 1996; Cawood et al., 2007; Li et al., 2008; see Slagstad et al., 2013 for an alternative model with later discussion in Möller et al., 2013) and the subsequent gravitational collapse. In addition, they possibly have pre-Sveconorwegian histories and some of them were reactivated also after the Sveconorwegian orogeny. The boundary zone between the Kongsberg and Telemark lithotectonic units of southern Norway (Fig. 1b), which is the subject of this contribution, is traditionally regarded as one of these major shear zones and has been shown on maps as a first order terrane boundary (Dons and Jorde, 1987; Nordgulen, 1999; Bingen et al., 2008a). The Kongsberg-Telemark Boundary Zone (from here on referred to as KTBZ) has already been previously reported as having accommodated several episodes of both ductile and brittle deformation (e.g., Starmer, 1979, 1985;Figs. 1 and 2). However, the boundary between the Kongsberg and Telemark lithotectonic units has in general been associated only with a major brittle fault zone named “Store Rivningsbreksje” (SR) by Bugge (1928), literally meaning “Great Friction Breccia”. This traditionally drawn boundary fades northward before terminating into a “triple junction” (Fig. 1b). This has caused confusion among researchers and the lack of a well-constrained structural and temporal evolutionary model for the KTBZ still prevents a refined understanding of the tectonic evolution of this sector of the orogen.

In this study, we aim to bridge this gap by presenting the results of a comprehensive structural investigation laying emphasis on the Sveconorwegian evolution of the KTBZ, in order to provide much needed structural and chronological constraints. We first present field observations from several transects across the boundary zone, and describe fabrics, fault rocks and kinematic indicators from either side. We also present a regional-scale foliation analysis from the northern part of the study area, where ductile geometries become more complex and bear significant implications for the regional geological picture. Microtectonic and EBSD (Electron Backscatter Diffraction) analysis of microstructures from gneisses and fault rocks are presented and discussed in order to support the structural (field) interpretation. In addition, we report U–Pb zircon geochronology from orthogneiss across the KTBZ to assess the age and origin of the protolith hosting the studied ductile fabrics. Our new results are then integrated to propose a model for the tectonic history recorded along the KTBZ. We discuss the implications of the new results for an improved understanding of the dynamics of strain accommodation along long-lived Sveconorwegian deformation zones and along Precambrian orogen-scale deformation zones in general. 2. Regional geological setting The Kongsberg lithotectonic unit in southern Norway is one of several lithotectonic units of the Sveconorwegian orogen. It was accreted and reworked during the Mesoproterozoic to Neoproterozoic Grenvillian-Sveconorwegian orogeny (1.14–0.90 Ga; Bingen et al., 2008b; Bogdanova et al., 2008). It is bound by the Telemark lithotectonic unit in the west and the Idefjorden lithotectonic unit in the north and is cut by the Permian Oslo Rift in the east (Fig. 1). The Kongsberg lithotectonic unit abuts the Telemark lithotectonic unit along the KTBZ, a deformation zone hosting various generations of both ductile and brittle fault rocks including those of the SR (Bugge, 1928, 1937). The latter represents a conspicuous

80

T. Scheiber et al. / Precambrian Research 265 (2015) 78–103

Fig. 2. (a) Example of airborne magnetic data used in the study (first vertical derivative of the total magnetic field) draped on the hillshaded Digital Elevation Model (DEM) of the Kongsberg lithotectonic unit. Traces of previously defined major structural elements are shown as: (1) “Store Rivningsbreksje” (Bugge, 1928); (2) Saggrenda-Prestfoss, (3) Prestfoss-Sokna and (4) Hokksund-Solumsmo mylonite zones (Starmer, 1985); (2) + (3) Saggrenda-Sokna shear zone (Bingen et al., 2001); (5) Vardefjell shear zone (Bingen et al., 2008a); (6) part of the Åmot-Vardefjell shear zone (Bingen et al., 2001); (1) + (5) part of the Kristiansand-Bang shear zone after Hageskov (1980) and Kongsberg-Telemark boundary (Andersen et al., 2002). (b) Combined lithotectonic and structural map of the Kongsberg lithotectonic unit and adjacent units.

c. 100 km long brittle fault zone, striking NNW–SSE in its southernmost part, N–S in the center and NE–SW in its northernmost extension (Figs. 1 and 2). In the south, it is cut by the Oslo Rift boundary faults (Rohr-Torp, 1973) whereas in the north it becomes progressively less prominent and it eventually fades away north of the village of Sokna (Fig. 2). Ductile precursors to the brittle SR were recognized along parts of the SR already by Starmer (1979). These were named according to their geographic location and are referred to as the Saggrenda-Prestfoss mylonite zone (SPMZ, zone 2 in Fig. 2a) and Prestfoss-Sokna mylonite zone (PSMZ, zone 3 in Fig. 2a; Starmer, 1985), the Saggrenda-Sokna shear zone (zone 1 + 2 in Fig. 2a; Bingen et al., 2001) and more generally the KongsbergTelemark boundary (zone 1 + 5 in Fig. 2a; Andersen et al., 2002). Starmer (1985) recognized a mylonitic shear zone also within the

Kongsberg lithotectonic unit, along the western side of the Modum metasedimentary complex and termed it the Hokksund-Solumsmo shear zone (zone 4 in Fig. 2a). The definition of the northern limit of the Kongsberg lithotectonic unit remains a matter of discussion. Starmer (1985) discontinued the boundary at the northeastern tip of the SR. Bingen et al. (2001) suggested that the Kongsberg lithotectonic unit wedges out toward the north where it is bounded by the ÅmotVardefjell shear zone (zone 5 + 6 in Fig. 2a). Domains north and east of this shear zone were correlated with the Idefjorden lithotectonic unit east of the Oslo Rift (Figs. 1 and 2). Alternatively, Nordgulen (1999) and Andersen et al. (2002) proposed that Kongsberg rocks continue toward the northwest, north of the Vardefjell shear zone (Figs. 1 and 2b).

T. Scheiber et al. / Precambrian Research 265 (2015) 78–103

The Kongsberg lithotectonic unit consists of a heterogeneous assemblage of amphibolite- to granulite facies para- and orthogneisses (see Bingen et al., 2008a), with granulite-facies rocks being more common in the north. The different lithologies can be grouped into three main units (Fig. 2b): (1) The Kongsberg complex (appearing in a western and an eastern belt on either side of (2)) makes up the biggest part of the Kongsberg lithotectonic unit and is composed of biotite–hornblende bearing gneissic metasedimentary rocks and orthogneiss bodies of various size (Fig. 2b). The granodioritic to tonalitic gneisses have an age of c. 1530 Ma (Andersen et al., 2004a) but younger (c. 1500 Ma) granitic orthogneisses are also present (Bingen et al., 2005). (2) The Modum complex mainly consists of supracrustal quartz-rich lithologies, typically quartzite, mica schist, and sillimanite-rich gneiss. It also hosts minor orthoamphibole-cordierite rocks. Detrital zircon age spectra from the quartzite indicate a maximum depositional age of ∼1475 Ma, which suggests correlation with equivalent rock types in the Bamble lithotectonic unit (Bingen et al., 2001; Andersen et al., 2002). Metagabbros hosted by the Modum complex have been dated at c. 1200 Ma (Munz et al., 1994). (3) The Veme complex is exposed to the NE of the Kongsberg complex. It consists mainly of mica schists and quartz-rich metasedimentary rocks originating from greywacke, quartz- or feldspar-rich sandstone and argillite. It has been correlated with the Stora Le-Marstrand formation east of the Oslo Rift (Bingen et al., 2001). Sedimentary structures such as graded bedding are locally preserved. Detrital zircon data from one sample from a turbidite facies within this unit yields a maximum depositional age of ∼1530 Ma (Bingen et al., 2001). The Veme complex hosts also meta-igneous bodies, two of them dated at c. 1555 Ma (Bingen et al., 2005) (Follum metadiorite pluton in Fig. 2b) and c. 1495 Ma (Bingen et al., 2008a). The KTBZ is localized within and along the eastern margin of a large N–S trending variably deformed granitic belt (Fig. 2b). One sample from this belt was dated at 1146 ± 5 Ma (Bingen et al., 2003), a time period marking the onset of the Sveconorwegian orogenic evolution. The c. 930 Ma old Flå granite (Nordgulen, 1999) in the north (Fig. 2b) cuts discordantly the structural fabric of the Kongsberg rocks, such that the formation of the main Sveconorwegian ductile fabric in the area is bracketed by these two intrusion episodes. The timing of the regional peak metamorphism was initially constrained for the Kongsberg complex at 1.2–1.1 Ga by Rb–Sr dating (Jacobsen and Heier, 1978). More recent metamorphic zircon, monazite, titanite and molybdenite dates from the Modum complex narrow the main metamorphic event down to c. 1110–1080 Ma (Bingen et al., 2008a and references therein). In addition to this metamorphic event, younger monazite and titanite ages (c. 1050–1025 Ma) have been reported from the Veme complex (Bingen et al., 2008a). The central part of the Telemark lithotectonic unit comprises a low-grade (greenschist- to lower amphibolite facies metamorphic conditions) supracrustal sequence, where volcanic and sedimentary structures and stratigraphic relations are widely preserved (Dons, 1960; Laajoki et al., 2002; Köykkä and Lamminen, 2011). Supracrustal rocks were deposited between ∼1510 Ma (Rjukan Group) and ∼1100 Ma (Eidsborg Formation) and are separated by several unconformities (Laajoki et al., 2002; Bingen et al., 2005; Lamminen and Köykkä, 2010; Spencer et al., 2014). Andersen et al. (2004b) suggested a depositional age of ∼1155–1145 Ma for the Norefjell quartzite (Fig. 2b). The Telemark lithotectonic unit mostly lacks the c. 1550–1530 Ma granodioritic to tonalitic orthogneiss typical of the Kongsberg lithotectonic unit (Andersen et al., 2004a), but hosts instead voluminous c. 1510–1500 Ma old volcanic rocks in its central part (Sigmond, 1998; Pedersen et al., 2009). Toward the margin of the Telemark lithotectonic unit there is an increase in metamorphic grade. Immediately to the west of the SR, our new data point to amphibolite-facies conditions. Similarly toward the

81

north, in the Hallingdal complex (Fig. 2b), amphibolite-facies metamorphism is dated at 1014 ± 1 Ma (metamorphic zircon, Bingen et al., 2008a). This metamorphism slightly pre-dates the activity of the Vardefjell shear zone (zone 5 in Fig. 2a), which is described as a SW-dipping sinistral shear zone, dated at 1012 ± 7 (zircon) to 985 ± 5 Ma (titanite; Bingen et al., 2008a). 3. Methods 3.1. Field data Field data were collected in synergy with the “KONGMO” mapping project by the Geological Survey of Norway (NGU), wherein the Kongsberg and Modum complexes were remapped in combination with new scientific investigations aimed at better constraining the tectonic evolution for this sector of the Sveconorwegian orogen. The simplified geological map of Fig. 2b is based on more than 20,000 observation points and reflects the new map, which is being currently finalized. Mapping benefitted greatly from modern airborne geophysical data flown and processed by the NGU. The new geophysical data include total magnetic field and full spectrum radiometric measurements. The data were acquired with 200 m flight line spacing along a flying direction of 140◦ (NW–SE) at a height of c. 60 m above ground. The interpretation of the geophysics for the purpose of this paper is largely qualitative, based on visual examination of images of the data. 3.2. Electron backscatter diffraction Four oriented polished thin sections were additionally polished with colloidal silica in order to reduce damages at the sample surface. Analyses were carried out on a ZEISS Evo 50 Scanning Electron Microscope (SEM), equipped with an EBSD camera (Adams et al., 1993), under low-vacuum conditions (10–15 Pa) and 20 nA beam current at the University of Bern. Grain orientation maps and pole figures of crystal directions based on these maps were obtained and calculated with the OIM® software package. Step sizes of 5 ␮m were chosen in order to eliminate artificial single-pixel measurements. Raw data were processed following standard OIM® data cleanup procedure. 3.3. U–Pb geochronology Zircon was separated from six crushed granitoid samples, using a water table, heavy liquids and magnetic separation. Selected hand-picked zircon crystals were mounted in epoxy together with chips of the reference zircon, and polished to approximately half thickness. The grains were imaged individually with a panchromatic cathodoluminescence (CL) detector in a variable pressure SEM. U–Pb analyses were performed by Secondary Ion Mass Spectrometry (SIMS) with the Cameca IMS 1280 instrument at the NORDSIM laboratory in Stockholm. The primary oxygen beam is c. 15 ␮m in diameter. Analytical protocols and data reduction follow Whitehouse et al. (1999) and Whitehouse and Kamber (2005). Analyses were calibrated using the 91500 Geostandard reference zircon (1065 Ma, Wiedenbeck et al., 1995), measured at regular intervals. The analyses were corrected for common Pb using the 204 Pb signal, if this signal was above background. SIMS data are reported in Table 1. Concordia age calculations (Ludwig, 1998), 207 Pb/206 Pb weighted average age calculation and TeraWasserburg (inverse) concordia diagrams were prepared with the ISOPLOT macro for Microsoft Excel (Ludwig, 2008). Uncertainties on the concordia and 207 Pb/206 Pb ages are quoted at 2 sigma (decay constant uncertainties propagated; systematic uncertainties resulting from interlaboratory experiments not propagated).

82

Table 1 SIMS zircon U–Pb geochronological data, corrected for common Pb. Sample/spot #

(a)

Zoning

(b)

Pit [␮m]

U [ppm]

Th [ppm]

Pb [ppm]

206

Pb Pb measured 204

(c)

f206 [%]

207 235

Pb U

sigma [%]

206 238

Pb U

sigma [%]

(d)

238 206

U Pb

sigma [%]

207 206

Pb Pb

sigma [%]

207

Pb Pb [Ma]

sigma

206

206

Pb U [Ma]

sigma

238

1.0 1.3 1.2 1.2 1.4 1.1 1.3 1.2 1.3 1.1 1.4 1.1 1.3 1.2

0.88 0.87 0.83 0.86 0.90 0.83 0.79 0.92 0.93 0.84 0.93 0.83 0.92 0.84

-3.7 -1.3 -0.8 -0.5 0.1 0.6 0.7 0.8 1.7 2.0 2.1 2.8 3.5 6.5

5.329 5.143 5.104 5.076 5.148 5.020 5.109 5.066 5.038 5.041 5.033 4.996 4.972 5.058

1.0 1.3 1.2 1.2 1.4 1.1 1.3 1.2 1.3 1.1 1.4 1.1 1.3 1.2

0.07805 0.07849 0.07860 0.07870 0.07786 0.07870 0.07792 0.07825 0.07811 0.07796 0.07796 0.07800 0.07791 0.07609

0.5 0.7 0.8 0.7 0.7 0.7 1.0 0.5 0.5 0.7 0.5 0.8 0.6 0.8

1148 1159 1162 1165 1143 1165 1145 1153 1150 1146 1146 1147 1145 1098

11 14 15 15 13 15 21 10 11 13 11 15 11 16

1109 1145 1153 1159 1144 1171 1152 1161 1167 1167 1168 1176 1181 1163

10 13 13 13 15 12 14 13 14 11 15 12 14 13

0.1947 0.1991 0.1996 0.2025 0.2017

1.5 1.2 1.3 1.3 1.3

0.94 0.62 0.94 0.91 0.89

-2.1 -2.0 0.4 2.3 2.6

5.136 5.023 5.009 4.938 4.957

1.5 1.2 1.3 1.3 1.3

0.07887 0.07980 0.07890 0.07869 0.07839

0.6 1.5 0.5 0.6 0.7

1169 1192 1170 1164 1157

11 30 9 12 13

1147 1170 1173 1189 1185

16 13 14 14 14

1.4 1.4 1.4 1.4 1.3 1.3 1.3 1.2 1.4 1.2 1.3 1.4 1.3 1.2 1.2

0.1897 0.1952 0.1940 0.1948 0.1983 0.1995 0.1980 0.1993 0.1969 0.1981 0.1992 0.1989 0.1997 0.2002 0.2031

1.2 1.3 1.2 1.3 1.2 1.2 1.3 1.1 1.3 1.1 1.3 1.3 1.2 1.2 1.0

0.88 0.92 0.86 0.92 0.91 0.93 0.96 0.90 0.93 0.94 0.96 0.94 0.93 0.96 0.89

-5.4 -3.9 -2.6 -2.6 -1.9 -1.1 -1.0 -0.8 -0.7 -0.6 0.6 0.8 1.0 1.2 3.3

5.272 5.124 5.155 5.133 5.042 5.011 5.051 5.017 5.080 5.048 5.019 5.027 5.008 4.996 4.923

1.2 1.3 1.2 1.3 1.2 1.2 1.3 1.1 1.3 1.1 1.3 1.3 1.2 1.2 1.0

0.07924 0.07978 0.07896 0.07912 0.07959 0.07953 0.07912 0.07932 0.07873 0.07897 0.07871 0.07855 0.07863 0.07864 0.07843

0.6 0.5 0.7 0.5 0.5 0.5 0.4 0.5 0.5 0.4 0.3 0.5 0.5 0.3 0.5

1178 1192 1171 1175 1187 1185 1175 1180 1165 1171 1165 1161 1163 1163 1158

13 11 14 11 10 9 8 11 10 8 7 9 9 7 11

1120 1149 1143 1147 1166 1173 1164 1172 1158 1165 1171 1170 1174 1176 1192

12 13 13 13 13 13 14 12 13 12 14 14 12 13 11

1.7 1.7 1.4 1.3 1.6 1.4 1.3 1.2 1.3 1.2

0.1866 0.1957 0.1988 0.1976 0.1939 0.1995 0.2020 0.2061 0.2078 0.2079

1.5 1.4 1.3 1.2 1.6 1.3 1.3 1.2 1.3 1.1

0.89 0.85 0.96 0.97 0.98 0.94 0.97 0.97 0.94 0.95

-5.5 -3.3 -2.0 -1.4 -1.2 -0.1 2.7 3.6 4.4 5.0

5.358 5.111 5.031 5.062 5.157 5.012 4.950 4.852 4.812 4.810

1.5 1.4 1.3 1.2 1.6 1.3 1.3 1.2 1.3 1.1

0.07860 0.07964 0.07975 0.07922 0.07832 0.07907 0.07844 0.07888 0.07894 0.07869

0.8 0.9 0.4 0.3 0.3 0.5 0.3 0.3 0.4 0.4

1162 1188 1191 1178 1155 1174 1158 1169 1171 1164

15 18 7 6 7 9 7 6 9 7

1103 1152 1169 1162 1143 1173 1186 1208 1217 1218

15 15 14 13 17 14 14 13 14 12

0.1876 0.1944 0.1959 0.1970 0.1942 0.1992 0.1957 0.1974 0.1985 0.1984 0.1987 0.2002 0.2011 0.1977

GVI25 (n4027, NGU78529) protoclastic augen gneiss, x: 525,763 − y: 6,625,712 (g) os 23,fo 291 90 66 18,914 0.10 K2-01a(S) K2-04a(S) os 23 34 14 8 >1e6 {0.00} K2-02a(S) un 23 370 60 83 >1e6 {0.00} 79 67 >1e6 {0.00} 283 23,fo os K2-03a(S) os 23,in 238 80 56 27403 0.07 K2-05a(S)

2.1171 2.1907 2.1715 2.1973 2.1804

1.6 2.0 1.3 1.4 1.4

GVI26 (n4018, NGU78509) strongly deformed granitic gneiss, x: 525,188 − y: 6,625,808 (g) os 22,fo 287 114 65 5309 0.35 2.0724 K1-11a(N) os 22 192 70 45 >1e6 {0.00} 2.1469 K2-02b(S) os 22 182 49 41 163,266 {0.01} 2.1121 K2-01b8S) {0.01} 2.1251 194,269 54 83 231 22 os K2-03b(S) os 22 284 158 70 377,536 {0.00} 2.1767 K1-04a(S) os 22 262 91 62 389,155 {0.00} 2.1881 K1-05a(S) 22 357 104 83 109,130 0.02 2.1598 os K2-03a(S) os 22 191 68 45 153,383 {0.01} 2.1797 K1-08a(S) os 22 226 77 53 114,134 0.02 2.1371 K2-02a(S) os 22 377 119 88 146,048 {0.01} 2.1571 K1-02a(S) 124 470,124 {0.00} 2.1623 116 545 22 os K2-04a(S) K2-01a(S) os 22,in 238 89 57 76,613 0.02 2.1545 os 22 254 97 61 68,836 0.03 2.1648 K1-03a(S) os 22 455 138 107 235,123 {0.01} 2.1705 K1-07a(S) {0.01} 2.1965 183,060 76 187 292 22 os K1-13a(S) BBI02 (n4011, NGU27052) foliated granite, x: 524,051 − y: 6,629,291 (g) D,os 18,po 1138 311 244 5545 K1-07a(N) co 18 152 36 34 33,467 K1-01a(S) D,U 18 682 228 160 401,581 K1-02a(S) D,os 18 1110 350 257 221821 K1-03a(S) 281 4086 270 20 1271 D,os K1-07b(S) K1-06a(S) D,os 18 398 110 92 >1e6 D,U 20 948 295 224 81,766 K1-13a(S) 20 919 235 219 170,133 D,U K1-11a(N) 154,135 161 222 654 D,os 18,fo K1-05a(N) os 20 604 203 148 357,451 K1-12a(N)

T. Scheiber et al. / Precambrian Research 265 (2015) 78–103

(f)

1.2 1.5 1.4 1.4 1.6 1.3 1.7 1.3 1.4 1.3 1.5 1.4 1.4 1.5

2.0226 2.1485 2.1857 2.1580 2.0943 2.1750 2.1848 2.2413 2.2620 2.2558

Disk. conv. [%]

(e)

BBI09 (n4024, NGU77309) coarse foliated granite, Bjørndalsetra, x: 533,799 − y: 6,602,341 (g) os 23 603 223 134 29,158 0.06 2.0192 K2-05a(N) 0.33 2.1041 68 5752 299 97 23,fi os K2-07a(S) os 23 142 68 34 87,136 {0.02} 2.1234 K2-04a(S) {0.03} 2.1378 61,958 46 36 23 156 os K2-12a(S) K2-18a(S) os 23 226 76 51 28,283 0.07 2.0853 58,271 {0.03} 2.1618 K2-02a(S) 63 44 187 os 23 K2-05b(S) os 23 158 44 36 16,322 0.11 2.1028 0.03 2.1299 77 56,528 330 115 os 23 K2-06a(S) os 23 306 102 71 83,106 0.02 2.1375 K2-08a(S) os 23 224 90 53 45,111 0.04 2.1322 K2-02b(S) 72 52,400 0.04 2.1359 307 103 os 23 K2-14a(S) os 23 155 55 37 38,132 0.05 2.1527 K2-06b(S) 58,267 0.03 2.1608 K2-17a(S) 112 70 290 os 23 os 23 144 68 35 22,700 0.08 2.0744 K2-08b(N)

0.34 {0.06} {0.00} {0.01} 0.46 {0.00} 0.02 {0.01} {0.01} {0.01}

rho

Table 1 (Continued) Sample/spot #

(a)

Zoning

(b)

Pit [␮m] (c)

U [ppm]

Th [ppm]

Pb [ppm]

206

Pb Pb measured 204

f206 [%]

207 235

Pb U

sigma [%]

206 238

Pb U

sigma [%]

(d) 2.0983 2.1065 2.1408 2.1112 2.1422 2.0854 2.1653 2.0829 2.1774 2.1261 2.1299 2.1670 2.2085

2.3 1.3 1.3 1.6 1.3 1.5 1.5 1.3 1.7 1.5 1.4 1.3 1.3

0.1907 0.1914 0.1946 0.1935 0.1957 0.1928 0.1978 0.1930 0.1991 0.1962 0.1982 0.2009 0.2043

BBI42 (n4016, NGU77334) feldspar porphyry, Prestfoss, x: 534597 − y: 6,656,360 (g) os 22 71 26 17 40,505 {0.05} K1-02a(N) os 22,fi 66 33 16 >1e6 {0.00} K1-08a(S) os 22 185 57 42 61,930 {0.03} K1-07a(S) os 22 195 102 47 136,994 {0.01} K1-09a(S) K2-10a(S) os 22 128 37 29 >1e6 {0.00} os 22 80 27 19 42,229 {0.04} K2-01a(S) D,os 22 520 144 119 206,620 {0.01} K1-01a(S) os 22 209 149 54 41,494 {0.05} K1-05a(S) se 22 123 27 28 >1e6 {0.00} K2-02a(S) os 22 118 71 29 30,399 0.06 K1-04a(S) os 22 226 102 55 178,939 {0.01} K1-09b(S) os 22 58 28 14 47,240 {0.04} K1-04b(S) os 22,fi 149 39 35 32,881 0.06 K2-09a(S) os 22,fi 261 76 59 20,555 0.09 K1-06a(N)

2.2287 2.1292 2.1018 2.1707 2.1378 2.1621 2.1462 2.1751 2.1865 2.1471 2.1511 2.1297 2.1457 2.0792

1.7 1.7 1.7 1.2 1.4 1.5 1.2 1.5 1.8 1.5 1.3 2.1 1.5 1.4

0.1996 0.1929 0.1916 0.1966 0.1949 0.1966 0.1966 0.1993 0.2003 0.1980 0.1988 0.1984 0.1995 0.1942

Disk. conv. [%]

238 206

U Pb

sigma [%]

207 206

Pb Pb

sigma [%]

207

Pb Pb [Ma]

sigma

206

206

Pb U [Ma]

sigma

238

(e)

(f)

1.7 1.1 1.2 1.4 1.2 1.4 1.3 1.1 1.5 1.2 1.2 1.2 1.1

0.73 0.88 0.90 0.89 0.90 0.91 0.89 0.88 0.88 0.78 0.87 0.94 0.91

-6.1 -5.8 -4.2 -3.3 -2.8 -2.0 -1.8 -1.5 -0.9 -0.7 1.9 2.5 3.8

5.244 5.224 5.139 5.169 5.110 5.185 5.057 5.180 5.023 5.097 5.046 4.979 4.896

1.7 1.1 1.2 1.4 1.2 1.4 1.3 1.1 1.5 1.2 1.2 1.2 1.1

0.07980 0.07982 0.07979 0.07915 0.07940 0.07843 0.07941 0.07826 0.07932 0.07860 0.07795 0.07825 0.07842

1.6 0.6 0.6 0.8 0.6 0.6 0.7 0.6 0.8 1.0 0.7 0.4 0.5

1192 1192 1192 1176 1182 1158 1182 1153 1180 1162 1146 1153 1158

31 12 11 15 11 13 14 12 16 19 14 8 10

1125 1129 1146 1140 1152 1137 1163 1138 1170 1155 1165 1180 1198

17 12 13 15 13 15 14 12 16 13 13 13 13

1.5 1.3 1.5 1.1 1.2 1.3 1.2 1.3 1.7 1.3 1.2 1.9 1.3 1.3

0.85 0.75 0.87 0.89 0.87 0.85 0.93 0.93 0.93 0.86 0.91 0.87 0.83 0.93

-4.3 -5.6 -5.1 -3.8 -3.5 -3.1 -1.8 -0.5 0.1 0.1 1.0 2.2 2.4 0.6

5.011 5.184 5.219 5.085 5.131 5.086 5.087 5.018 4.992 5.051 5.029 5.041 5.013 5.150

1.5 1.3 1.5 1.1 1.2 1.3 1.2 1.3 1.7 1.3 1.2 1.9 1.3 1.3

0.08099 0.08005 0.07955 0.08006 0.07956 0.07976 0.07918 0.07917 0.07916 0.07866 0.07846 0.07787 0.07802 0.07766

0.9 1.1 0.9 0.5 0.7 0.8 0.4 0.5 0.7 0.7 0.6 1.1 0.8 0.5

1221 1198 1186 1198 1186 1191 1177 1176 1176 1163 1159 1144 1147 1138

18 21 17 11 13 16 9 11 13 15 11 21 17 10

1173 1137 1130 1157 1148 1157 1157 1171 1177 1164 1169 1166 1172 1144

16 13 16 12 12 14 12 14 18 14 13 20 14 14

T. Scheiber et al. / Precambrian Research 265 (2015) 78–103

GVI23 (n4017, NGU78532) protocataclastic augen gneiss, x: 526,088 − y: 6,647,675 (g) K1-02a(N) os 22,fi 82 45 19 13,442 0.14 os 22 152 51 34 91,237 {0.02} K1-01a(N) os 22 169 58 39 62,661 {0.03} K1-05b(N) os 22 122 47 28 81,153 {0.02} K2-05b(S) os 22 184 68 43 177,823 {0.01} K1-05a(S) os 22,fo,in 153 36 34 30,041 0.06 K2-02b(S) os 22 133 49 31 152,752 {0.01} K1-10a(S) os 22 247 86 56 112,779 {0.02} K1-03a(S) os 22 114 42 27 298,798 {0.01} K2-05a(S) K1-09a(S) os 22 62 29 15 53,453 {0.03} K2-01a(S) os 22 119 31 27 59,543 0.03 os 22 305 103 73 238,342 {0.01} K2-02a(S) K2-06a(S) os 22 214 62 51 70,196 0.03

rho

(a) Analyses selected for concordia age calculation are noted (S) and non-selected are noted (N). (b) Zoning: zoning defined from CL image: C: core; R: rim; D: CL dark; B: CL bright; un: unzoned-weak zoning; os: oscillatory growth zoning; se: sector zoning; co: convoluted zoning; ir: patchy-irregular zoning. (c) Pit: BSE image of analytical pit after analysis. Length of the pit in micron; fo: fracture extending inside and outside of the pit, fi: fracture limited to the inside of the pit, in: inclusion, po: porosity. (d) Proportion in % of 206 Pb of common Pb derivation, value between bracket indicates that 204 Pb is at background level and that no common Pb correction was performed. (e) Error correlation in conventional concordia space. (f) Age discordance in conventional concordia space. Positive numbers are reverse discordant. (g) X-Y UTM coordinates, WGS84, zone 32, collected by means of GPS.

83

84

T. Scheiber et al. / Precambrian Research 265 (2015) 78–103

4. Field observations along selected transects across the SR In order to unravel the origin and structural evolution of the KTBZ and place it in the broader framework of the Sveconorwegian history, we investigated selected areas across the traditionally drawn boundary zone, i.e., the SR (Fig. 2b). The trace of the SR as defined by Bugge (1928, 1937) was confirmed by our own observations. The SR is used for each section as a convenient marker lineament to separate the structural analysis between a western and an eastern block. This is because there are local, remarkable differences in the ductile structural patterns east and west of the SR.

associated with sinistral asymmetric clasts. The microstructure of this mylonitic rock is described in more detail in Section 6. Both the mylonite and the amphibolitic gneiss are cut across at their contact by a dense network of epidote- and chlorite-coated brittle faults. Sets of en-echelon fractures at different scales indicate E–W extension. Small-scale quartz veins are very common in the area. In summary, the SR defines a boundary between two distinct structural domains along this transect, wherein mylonites revealing a component of sinistral shear occur only east of it, ENE-vergent kink folds only west of it, and the stretching lineation of the gneissic fabric has a slightly different orientation on both sides (Fig. 4b). All these structures were overprinted by later intense fracturing and veining.

4.1. Kolsjø section 4.3. Flesberg area The southernmost studied traverse remains entirely within a homogeneous, medium- to coarse-grained granitic gneiss (Figs. 2b, 3a). The gneiss contains a constant steeply ENE-dipping penetrative biotite–chlorite foliation both to the west and east of the trace of the SR (Fig. 3b). There is no obvious ductile strain gradient developed along the investigated transect. No clear stretching lineation was observed on the main ductile foliation. The SR cuts discordantly the gneiss with several visible decimetric fault zones. Discrete faults and fractures are concentrated predominantly in the immediate vicinity to the SR and can be interpreted as belonging to a conjugate set of steeply-dipping faults having generally the same orientation as the SR, and steeply-dipping fractures which strike perpendicular to it. Both the granitic gneiss and the brittle fault zones are in turn cut along the entire transect by quartz veins (with individual veins up to 25 m in thickness), which are also most abundant closest to the trace of the SR (Fig. 3a). Dense quartz vein networks can locally lead to the brecciation of the host rock (Fig. 3c). Interestingly, quartz veins in the eastern part of the section strike mainly ENE-WSW, whereas in the west they generally trend N–S (Fig. 3b), which possibly indicates different fracturing histories on either sides of the SR. In summary, this southernmost transect displays abundant quartz veins and brittle faults. No obvious ductile precursor to this confined fault zone is developed. 4.2. Buvannet section In this section, the belt of granitic gneiss contains interlayered slivers of banded amphibolites (Fig. 4a). Both rock types contain a moderately (E)NE-dipping gneissic foliation and a clear downdip stretching lineation. The orientation of the stretching lineation changes slightly along the section from a NE-plunge in the west to an E-plunge in the east (Fig. 4a and b). No clear kinematic indicators were observed associated with the described foliation and lineation. Two generations of folds affect the main foliation: The older folds are open and have axes dispersed in the north-east quadrant with axial planes dipping gently or moderately to the east (Fig. 4b), but without an axial planar cleavage developed. The younger folds are restricted to the area west of the SR. They are ENEvergent kink folds with subhorizontal NNW–SSE trending fold axes and axial planes dipping gently toward the west (Fig. 4c). The consistent vergence of these folds points to an overall top-to-the-ENE movement post-dating the main fabric development. Toward its western contact to the amphibolitic gneiss, the granitic gneiss passes to an up to 100 m wide zone of mylonites (Fig. 4a). These are characterized by pinkish, rounded, mostly feldspathic clasts speckling a fine-grained dark matrix (Fig. 4d). Dark sealed fractures abound in this transition zone (Fig. 4e). The mylonitic foliation dips moderately to steeply to the E and is thus slightly discordant to the foliation in the adjacent gneiss (Fig. 4b). The gently SSE-plunging lineation developed on this foliation is

The Flesberg area provides access to particularly well-preserved exposures of the KTBZ. The area is structurally complex with significant variations in the orientation of the main structural features both along and across strike of the KTBZ. For clarity sake, we use again the SR to subdivide the mapped region in a western and an eastern block. The granitic gneiss in the area is interlayered with banded gneiss which widens toward the north (Fig. 5a). The main foliation west of the SR has a much more constant orientation than the foliation east of the SR (Fig. 5b–d). South and immediately north of Flesberg, this penetrative amphibolite-facies foliation strikes N-S, and is associated with a strong and moderately NE-plunging lineation defined by amphibole and severely stretched quartz and feldspar crystals (Fig. 5a–d). This stretching lineation is particularly well-developed in the proximity to the SR. Axes of tight to isoclinal rootless folds are parallel to this stretching lineation (Fig. 6a) and locally there is evidence for top-to-the-NE transport. In contrast, farther north the general strike of the main amphibolite-facies foliation changes gradually to NW–SE with gentle dips, that is, it becomes oblique to the SR, which traditionally is used in that area as the c. N–S contact between Telemark and Kongsberg (Fig. 2b). There the main ductile foliation is associated with a subhorizontal NW–SE trending stretching lineation in the banded gneiss (Fig. 5a and b). Foliation-parallel transposed pegmatitic dykes are boudinaged parallel to this stretching lineation (Fig. 6b), which is perpendicular to the stretching lineation close to the SR farther south. The coarse- to medium-grained granitic gneiss to the east of the SR is also characterized by a penetrative foliation. However, the NE-plunging stretching lineation is no more prevalent. Instead, a weakly-developed ESE-plunging lineation defined by elongate quartz aggregates (Fig. 5d) and a well-developed SSE-plunging stretching lineation in high-strain zones represent the common linear fabric (Fig. 5b and c). Shear bands in gently dipping granitic gneiss indicate top-to-the-NW shearing (Fig. 6c). Several outcrops preserving mylonitic to ultramylonitic fabrics identical to those observed in the Buvannet area (Fig. 4d) also occur within the granitic gneiss throughout the Flesberg area (Fig. 5a). The mylonitic foliation dips in general moderately to the ENE, and a gently SSEplunging stretching lineation is developed on the fine-grained matrix (Fig. 5b–d). Associated shear sense criteria such as asymmetric clasts (Fig. 6d) indicate invariably sinistral strike-slip. The gneissic foliation east of the SR is affected by folding with variable fold axis orientations (Fig. 5b–d). In the Beinsvatnet area (Fig. 5c), for example, asymmetric folding of the main foliation around a subvertical axis affects the granitic gneiss (Fig. 6e and f). The associated axial plane is subvertical, strikes N–S and asymmetries indicate sinistral strike-slip in map view. In other localities the fold axes plunge to the SSE (Fig. 5) that is subparallel to the stretching lineation associated with the sinistral kinematics in the mylonites.

T. Scheiber et al. / Precambrian Research 265 (2015) 78–103

85

Fig. 3. Structural data from the Kolsjø area with (a) structural map and cross-section and (b) stereoplots (lower hemisphere, equal area). (c) Network of quartz veins affecting the granitic gneiss (530425/6601326), coin for scale.

Ductile to brittle extensional structures affecting all the aforementioned features are common in the entire Flesberg area. They include (a) open and ENE-vergent folds with subhorizontal to gently ENE-dipping axial planes (Fig. 6g), which are particularly prevalent in the banded gneiss series west of the SR and point to an overall top-to-the-ENE transport; (b) conjugate sets of distinct steeply-dipping normal faults (Fig. 6h) indicating E-W extension and occurring on both sides of the SR, but concentrated in its immediate vicinity; (c) ductile foliation planes locally reactivated as lower-grade fault planes. Younger quartz veins, epidote-filled joints and brittle faults, often associated with cataclastic deformation (Fig. 6i) are also concentrated along the trace of the SR. In several outcrops quartz veins penetrate the host rock and are volumetrically so dominant that they represent the main rock type. Various generations of mutually cross-cutting veins can be seen, and brecciation of the host rock can be conspicuous, with local significant volumes of well-exposed dilational breccias (Fig. 6j). In summary, the area west of the SR in the Flesberg region is characterized by an intense moderately E- to NE-dipping ductile fabric, which is slightly oblique to the local trace of the SR. Top-to-the-NE shearing is dominant in the southern part, whereas a NW-SE-striking stretching lineation is present in the northern part. In contrast, the area east of the SR preserves a more variably oriented ductile fabric associated with top-to-the-(N)W shear deformation and contains several (ultra-) mylonitic lenses and folds, which invariably indicate sinistral kinematics. Late brittle extensional structures are prevalent close to the trace of the SR in the entire Flesberg area.

4.4. Prestfoss area In the southern part of this area (Lake Soneren, Fig. 7a), the main ductile fabric west of the SR is defined by a flat SE-dipping foliation which turns locally into a subvertical orientation (Fig. 7b). There

is a consistently SSE-plunging lineation, defined by quartz and feldspar mineral rods (LS-tectonites, Fig. 7c) and aligned amphibole needles, similar to that in the northern part of the Flesberg area (Fig. 5a). Shear senses, when observed, are inconsistent. In the northern part of the Prestfoss area the trace of the SR coincides with the lithological contact between granitic gneiss in the west and banded gneiss consisting of densely alternating granodioritic to dioritic orthogneiss and paragneiss in the east (Fig. 7a). There the area immediately west of the SR is characterized by a moderately E(SE)-dipping ductile foliation and a locally-developed stretching lineation associated with shear bands indicating top-to-the-ENE shearing (Fig. 7b and d). The area east of the SR is characterized by a subvertical generally N–S striking ductile fabric which flattens out toward the trace of the SR. Just west of Prestfoss SE-dipping foliation planes bear E-plunging stretching lineations, and vergent folds point to topto-the-W thrusting (Fig. 7e). The remaining observations show Sto SE-plunging stretching lineations on steeply-dipping foliation planes which are associated with sinistral shear fabrics (Fig. 7f and g), as similarly observed in the Flesberg area east of the SR. Tight to isoclinal folds affecting an existing foliation at various scales have fold axes parallel to this stretching lineation and subvertical axial planes. In addition, a lens of the same mylonitic lithology found in the Flesberg and Buvannet areas, where it is invariably associated with sinistral shear deformation, occurs within the banded gneiss series close to the SR (Fig. 7a). Later folding can be observed all over the area. Fold axes have similar orientations as the sinistral folds, but have different styles. West of the SR the main ductile fabric is folded openly into upright folds, while east of the SR micaceous lithologies are crenulated and locally associated with a gently SE-dipping crenulation cleavage. Younger brittle faults and quartz veins are again concentrated along the trace of the SR and have similar styles and orientations to what has been observed in the Flesberg area. In addition, a major brittle fault zone strikes perpendicular to the trace of the SR and

86

T. Scheiber et al. / Precambrian Research 265 (2015) 78–103

Fig. 4. Structural data from the Buvannet area. (a) Structural map and cross-section. (b) Stereoplots. (c) Kink folds affecting amphibolitic gneiss (526617/6610081), coin for scale. (d) Mylonite (526587/6610334), coin for scale. (e) Diffuse fracturing of the granitic gneiss (526678/6610387), coin for scale. Abbreviations: P, planes; L, lineations; FA, fold axis; AP, axial plane.

can be traced through the village of Prestfoss (Fig. 7a). An apparent sinistral offset can be recognized in map view. In summary, evidence for top-to-the-W thrusting is present locally just west of Prestfoss. Elsewhere, the area is characterized by a pervasive ductile fabric associated with sinistral shear along S(E)plunging lineations east of the SR, which is instead missing west of it. West of the SR, ductile NW–SE stretching is present in the southern part, whereas evidence for ductile top-to-the-NE extension is present in the northern part. Younger brittle structures and quartz veins define the trace of the SR in the entire area.

4.5. Sokna area In the area of Sokna, the trace of the SR is less clear and fades toward the NE (Fig. 2). However, the KTBZ is defined by similar structural characteristics as in the sections described above. Dioritic gneiss from the Kongsberg complex in the immediate east of the SR records localized shear zones at the meter-scale indicating top-to-the-W thrusting (Fig. 8a). Lenses of the characteristic mylonite occur west of the SR (Fig. 8b), where the main foliation strikes NW-SE (Fig. 2b). The stretching lineation plunges gently

T. Scheiber et al. / Precambrian Research 265 (2015) 78–103

87

Fig. 5. (a) Structural map of the Flesberg area. (b–d) Structural data provided in stereoplots and cross-sections for three transects. Orientation of (ultra-) mylonitic fabric is included in the stereoplots showing the ductile fabric E of the SR. Abbreviations: P, planes; L, lineations; FA, fold axis; AP, axial plane.

88

T. Scheiber et al. / Precambrian Research 265 (2015) 78–103

Fig. 6. Structural aspects from the Flesberg area: (a) Banded gneisses immediately to the west of the SR. Amphibole-defined stretching lineation (indicated by orange arrow) is parallel to fold axes (525187/6625815), coin for scale. (b) Banded gneiss W of the SR hosting boudinaged pegmatite (523119/6641628), coin for scale. Note that this stretch is orthogonal to the one documented in (a). (c) Shear bands in granitic gneiss east of the SR (526005/6630049), coin for scale. (d) Mylonitic fault rock with sinistral asymmetric clasts (526257/6636933), coin for scale. (e) Lineation associated with (f) asymmetric folding of the main ductile fabric around a steeply N-plunging fold axis (526937/6630827), hammer for scale. (g) ENE-vergent folding of the banded gneisses west of the SR (525369/6635185), magnifier for scale. (h) Normal faults indicating E-W extension (525187/6625815), coin for scale. (i) Fault-bounded cataclasite (526003/6630341), coin for scale. (j) Dilation and breccia formation by quartz veining (526154/6637669), hammer for scale.

T. Scheiber et al. / Precambrian Research 265 (2015) 78–103

89

Fig. 7. (a) Structural map of the Prestfoss area. (b) Structural data from different subareas (P, planes; L, lineations; FA, fold axis; AP, axial plane.). (c) LS-tectonite west of the SR with pronounced mineral rodding (530996/6655454), finger for scale. (d) Top-to-the-ENE shear bands indicating normal faulting close to the trace of the SR (536195/6662534), magnifier for scale. (e)–(g) Gneisses east of the SR. (e) Imbricated felsic layer within porphyritic layers of the banded gneiss series indicating top-tothe-W thrusting (534602/6656358), scale bar is 10 cm long. One sample from a porphyritic layer in this outcrop is dated at 1166 ± 6 Ma (BBI-42). In order to present all photographs in the same orientation this photograph is a mirror image. (f) Sigma clast (538687/6659937) and (g) delta clast (539233/6655978) indicating sinistral shear, coins for scale.

90

T. Scheiber et al. / Precambrian Research 265 (2015) 78–103

Fig. 8. Field aspects from the Sokna region. (a) Photograph and geological interpretation of apparently monophase deformed diorite (light gray) and pegmatitic dykes (red). Meter-scale S-C fabric indicates top-to-the-W shearing (553839/6679941). G. Viola for scale. (b) Mylonite NW of Sokna (547728/6680814). Typically pink rounded feldspathic clasts are bleached on this surface, coin for scale. (c) E-vergent folds with subhorizontal axial planes in migmatitic gneisses close to the trace of the SR (539942/6669942), hammer for scale. (d) Roadcut west of Sokna (543387/6681592). Gently SSW-dipping main foliation is overprinted by steeply-dipping top-to-the-ESE normal faults (dashed in stereoplot). Mafic layers in granitic gneiss are asymmetrically boudinaged. Note pegmatitic injections along normal faults, person for scale. Planes in stereoplots correspond to the main foliation when continuous, and to axial planes and steep normal faults (dashed). Circles represent stretching lineations and triangles fold axes.

toward the ESE and asymmetric clasts point to sinistral shearing (Fig. 8b). In the area west of the SR gneisses and mylonites are affected by ductile extensional structures, which are locally associated with migmatites. These include E-vergent folds at different scales (Fig. 8c) and, particularly in the area west and northwest of Sokna, sets of localized steeply ESE-dipping normal faults. These faults are associated with felsic melt injections, especially where they deform competent mafic layers (Fig. 8d). Extensional structures in the Sokna area seem to be present in a broad zone and transport direction is toward the E to ESE. Brittle faults and up to several tens of meter wide quartz veins (see also Nordgulen, 1999) occur predominantly along the trace of the SR. 5. Regional-scale foliation trajectory pattern in the Norefjell-Hønefoss area Whereas the main ductile foliation in both the Kongsberg and Telemark lithotectonic units runs parallel to subparallel to the trace of the SR from Saggrenda in the south to almost Sokna in the north, the regional-scale foliation pattern in the Sokna area becomes much

more complex and the trace of the SR fades out (Figs. 2, 9a). In order to better understand the geometry of this area we conducted a regional-scale analysis of foliation traces based on c. 2,000 foliation measurements collected during the NGU mapping program (Fig. 9a). This analysis reveals the existence of a major, complex fold structure that extends between Norefjell and Hønefoss (Figs. 2b and 9a). Parts of this fold were already reported from the Sokna region (“Sokna fold” after Jakobsen, 1991) and from south of Hønefoss (Starmer, 1985). Foliation analysis reveals that both of the reported fold structures may form part of the same largescale isoclinal fold, which is geometrically sandwiched between two metamorphosed gabbroic bodies (Holleia and Follum bodies, Fig. 9a and b). The geometry of the Norefjell-Hønefoss fold is very complex and changes gradually along the axial plane trace (see domains 1–8 in Fig. 9a and b). The meta-arkoses of the Veme complex are intensely affected by parasitic folds on many different scales (size orders 1 cm to 100 m), whereas the other lithologies in other domains are mostly devoid of small-scale fold structures. The sudden change in orientation of the axial plane trace north of the village of Sokna is

T. Scheiber et al. / Precambrian Research 265 (2015) 78–103

91

Fig. 9. (a) Foliation trajectory map and stereoplots showing the variable orientation of structural elements of the Norefjell-Hønefoss fold in domains 1–8. The axial plane trace of an earlier fold generation is shown by the short dashed line, the trace of a later fold generation is shown by the long dashed line. Note the constant orientation of foliation in domains 9 and 11 for comparison. (b) Schematic 3D sketch of the Norefjell-Hønefoss fold. Fold axes are given in red. (c) Calculated axial planes and fold axes for domains 1–8 and 10 in the present orientation and after applied rotation. Rotation axis is represented by the calculated fold axis of domain 10. Angle of rotation is 90◦ for domains 1 and 3, and 120◦ for domain 2. Rotated structures are given in red. Abbreviations: n, number of measurements; FA, calculated fold axis; AP, calculated axial plane. (For interpretation of the references to color in text, the reader is referred to the web version of this article.)

an important feature with implications for the regional geological picture. Geometrically, this change can be explained by refolding around a subvertical fold axis, which is defined by gneissic units wrapping around the Holleia gabbroic body (domain 10 in Fig. 9a and b). Rotation of the western fold segments (domains 1–3 in

Fig. 9a) with suitable angles (i.e., corresponding to the change of the axial plane trace in map view) around this fold axis (domain 10 in Fig. 9a) results in a fold which has a steeply W-dipping axial plane and axes roughly distributed along a great circle with a concentration around a steep westerly plunge (Fig. 9c). However, it is

92

T. Scheiber et al. / Precambrian Research 265 (2015) 78–103

difficult to assess whether it is the eastern or the western geometry that represents the original one. 6. Microstructural analysis According to the field observations from the Flesberg area (Section 4.3), the gneisses east and west of the trace of the SR are characterized by different macroscopic fabrics with different stretching lineation orientations. Additionally, several lenses of a typical mylonite occur. In order to strengthen the structural interpretation derived from field observations, the mylonitic fault rock is described in greater detail and EBSD analyses are presented from four samples covering different fabric types including the mylonite. 6.1. Microstructural description of the mylonitic fabric Along the margins of the mylonitic lenses, variable transitions from the gneissic host rock into fully developed (ultra-) mylonites can be studied (Fig. 10a–c). The granitic gneiss is characterized by pink to white K-feldspar clasts and elongate monomineralic recrystallized quartz aggregates. These define the regional foliation together with elongate K-feldspar, white mica and/or greenish biotite and at times blurry layers of ilmenite (Fig. 10a, d–f). Plagioclase, garnet and allanite also occur, and epidote overgrows this fabric. At the transition zones to the mylonite, the granitic gneiss reveals protocataclastic deformation, especially in K-feldspar clasts, which induces grain-size reduction and disturbs the gneissic foliation (Fig. 10d and e). Fracture sets sometimes form at an angle to the gneissic foliation and define a fracture cleavage (Fig. 10d, e and g). White mica, greenish biotite and ilmenite grow in this new cleavage (Fig. 10g). Quartz aggregates and K-feldspars become progressively disrupted, mixed and the grain size decreases dramatically toward the higher-strain mylonite. However, some Kfeldspar clasts remain undeformed during this deformation process and provide the mylonite with its characteristic augen fabric. The clast-matrix ratio is variable, especially at transition zones. Sometimes a foliation-parallel layering can be recognized, where layers contain different amounts of clasts (Fig. 10h). The mylonite appears as a very hard fault rock, locally without a well-developed macroscopic foliation (Fig. 10b), elsewhere, with a strong mylonitic foliation (ultramylonite, Fig. 10c). Similar mylonitic fault rocks lacking a macroscopic foliation have been described elsewhere (e.g., Simpson et al., 2001). In thin sections, however, a foliation can always be recognized. This foliation bends around the clasts and is defined predominantly by finely dispersed greenish biotite. In (ultra-) mylonites K-feldspar clasts are extremely rounded and are intergrown with quartz and biotite especially at their sutured rims (Fig. 10i). Garnet and allanite occur in the fine-grained matrix composed of quartz, K-feldspar, plagioclase and biotite. Epidote is abundant in the mylonite.

the ductile fabric with a NE-plunging stretching lineation present within the KTBZ west of the SR from the Buvannet area in the south, through most of the Flesberg area and to the northern part of the Prestfoss area (Figs. 4, 5 and 7). The foliation is defined by chlorite, replacing biotite, and is concordant to an amphibole-defined foliation in adjacent mafic layers. Quartz and feldspar grains are coarse and equigranular (mean recrystallized quartz grain size of c. 300 ␮m) and mineral aggregates show a strong shape preferred orientation (Fig. 11a), which defines the NE-plunging stretching lineation at the outcrop. K-feldspar aggregates probably represent remnants of larger magmatic augen, which were later reworked by ductile shearing. Large quartz grains show weak undulose extinction. The pattern of the Crystallographic Preferred Orientation (CPO) of c-axes is dominated by a Y-axis maximum and subordinate clusters at high angle to the trace of the foliation in the periphery of the stereogram (Fig. 11a). TSC-13 is from a gneiss of granitic composition east of the SR. It contains white to pink, up to 3 cm long K-feldspar crystals, which are aligned according to the foliation. The orientation of the foliation is similar to that of sample TSC-56, but the weakly developed stretching lineation plunges toward the SE and not to the NE (Fig. 5a). This sample is furthermore characterized by monomineralic recrystallized quartz aggregates (up to 1 cm in diameter), which are also aligned along the foliation planes, but have irregular shapes and sinistral asymmetric tails (Fig. 11b). These aggregates host large equidimensional grains (mean recrystallized grain size of 235–260 ␮m) with mostly straight grain boundaries. The analyzed aggregates are characterized by a clear CPO with the c-axes forming a single Y-maximum and a noticeable concentration in the periphery of the stereogram at high angle to the trace of the foliation ((1) and (2) in Fig. 11b). TSC-28 is a sample from a mylonitic lens characterized by a clear foliation and by the presence of quartz ribbon grains and sub-rounded K-feldspar clasts in a fine-grained quartzofeldspathic matrix. Biotite and epidote occur as well. Quartz ribbons are locally observed to be partially resorbed, which leads to phase mixing and grain size reduction (Fig. 11c). As a consequence, the grain size distribution of quartz is bi- or even polymodal with a mean recrystallized grain size of around 86 ␮m. Quartz grains bear a clear Y-maximum of c-axes orientations (Fig. 11 c). TSC 62 represents a strongly foliated ultramylonite characterized by large rounded clasts speckling a fine-grained matrix (Figs. 6d, 10c and 11d), which consists predominantly of quartz, K-feldspar, plagioclase and biotite. Clasts are formed by large single K-feldspar crystals and are wrapped around by a biotite-defined foliation. Accessory minerals are epidote, garnet, titanite, allanite and zircon. Grains within the matrix are equigranular and recrystallized quartz grains are between 20 and 50 ␮m in diameter (Fig. 11d). Quartz grains hardly form interconnected layers and are randomly scattered in the fine matrix. Nonetheless, a weak CPO is shown in the pole figures (Fig. 11d).

6.2. EBSD textural analysis 6.3. Microstructural and -textural interpretation Four samples from the Flesberg area (Fig. 5a), representing different varieties of granitic gneiss and mylonitic fault rocks, were selected for detailed microstructural analysis. Texture analysis by means of EBSD was carried out on recrystallized quartz grains from mono- and polymineralic fabric domains (Fig. 11). Photographs and inverse pole figure maps of Fig. 11 illustrate sections perpendicular to the mylonitic foliation (fn ) and parallel to the stretching lineation (L) on the sample. TSC-56 is a strongly foliated and lineated granitic lens from within the banded gneiss series, sampled immediately west of the trace of the SR (Fig. 5a). Dating of zircon in sample GVI-26 from the same rock and same locality yields a magmatic protolith age of 1169 ± 5 Ma (see Section 7). This sample is representative for

The microstructural description of the mylonite occurrences and the EBSD analyses of four different fabrics allow conclusions about shear senses, deformation conditions and the deformation mechanisms, which were active during the different fabric-forming events. The quartz c- and a-axis patterns obtained from the four samples can be classified into two main texture types, i.e., Y-maxima and concentrations in the periphery of the c-axis pole figures, arranged on single girdles. These resemble the commonly observed quartz dislocation creep patterns described by Schmid and Casey (1986), which can be interpreted to result from dislocation glide on basal and/or prism slip systems within quartz (e.g., Law, 1990;

T. Scheiber et al. / Precambrian Research 265 (2015) 78–103

93

Fig. 10. Scanned polished rock samples showing (a) a granitic gneiss from east of the SR and its (b) mylonitic and (c) ultramylonitic equivalents. Note the lack of strong macroscopic foliation and preservation of rounded K-feldspar clasts in (b). The ultramylonitic sample (c) was selected for microstructural analysis (TSC-62). (d and e) Crosspolarized and plane-polarized light micrographs showing microstructures at the transition from granitic gneiss to mylonite. Protocataclastic deformation and associated fracture cleavage developed at an angle to the gneissic foliation, which is characterized by elongate quartz aggregates (f). Newly grown minerals are associated with the new cleavage (g). (f) and (g) are close-ups to (d). (h) Heterogeneous distribution of K-feldspar clasts, probably as a result of strain. (i) Highly rounded K-feldspar clasts in a finegrained matrix, typical for the ultramylonite. Clast rims are characterized by sutured grain boundaries with mainly quartz intergrowth. Green arrows in microphotographs indicate protocataclastic deformation of K-feldspar clasts, white arrows point out newly grown mineral phases.

94

T. Scheiber et al. / Precambrian Research 265 (2015) 78–103

Fig. 11. Thin section photographs (cross-polarized light), grain orientation maps (inverse pole figure color coding with reference direction = Z), contoured pole figures (c-axes and a-axes) and grain size distributions of dynamically recrystallized quartz from oriented samples TSC-56 (a), TSC-13 (b), TSC-28 (c) and TSC-62 (d). Samples are oriented parallel to the lineation (L) and perpendicular to the foliation (fn ). Sample TSC-28 has been rotated into parallelism with the lineation during data processing. However, note the different section orientation and the different scales of photographs and pole figure maps. The microstructures and textures from (b) to (d) highlight the progressive fabric evolution leading to the ultramylonite (d). This transition is characterized by progressive dispersion of quartz, decrease in mean grain size and CPO weakening. Logarithmic contouring of CPO intensities is indicated next to corresponding pole figures. Abbreviations: max = maximum pole density of c-axes; n = number of grains analyzed; GS = grain size (␮m); AF = area fraction (%); d = average grain size diameter.

T. Scheiber et al. / Precambrian Research 265 (2015) 78–103

Heilbronner and Tullis, 2006; Toy et al., 2008). In sample TSC56 (Fig. 11a), the c-axes girdle is rotated clockwise with respect to the X–Z plane, which corresponds to a geographic top-to-theNE shear sense. Similarly, weakly asymmetric single girdles can be recognized in sample TSC-13 (Fig. 11b), but they are inconsistent: (1) shows a very weak dextral asymmetry, whereas (2) has a more pronounced sinistral asymmetry. This sinistral asymmetry together with asymmetric clast tails indicate a top-to-the-NW sense of shear in geographic space. Analyzed quartz-rich layers from sample TSC-28 show a pronounced Y-maximum of quartz caxes, which is characteristic for simple shear deformation (Schmid and Casey, 1986) but does not allow an interpretation of the transport direction. The weak, though noticeable asymmetries measured in the strongly foliated mylonite (sample TSC-62) indicate a sinistral sense of shear, which agrees well with the sinistral kinematic indicators observed at the outcrop scale (Fig. 6d). However, since this rock most likely deformed by diffusion creep with grain boundary sliding mechanisms (e.g., Stünitz and Gerald, 1993; Paterson, 1995; Kilian et al., 2011), textures are difficult to interpret. Samples TSC-56 and TSC-13 are both from granitic gneiss and contain an ENE-dipping ductile foliation, but were collected from either sides of the SR (Fig. 5). Our textural and microstructural analysis shows that they reflect different deformation mechanisms. From monomineralic quartz layers of the gneissic sample TSC56, it can be seen that its microstructure reflects grain boundary migration (GBM) recrystallization, which is generally associated with high-temperature deformation at >500 ◦ C (e.g., Stipp et al., 2002), provided that grain boundary mobility is ensured. The coarse-grained equigranular microstructure, together with the monomineralic quartz and feldspar ribbons, which gives the rock a banded appearance, also points to high-temperature deformation (Hippertt et al., 2001). However, pseudomorphs of chlorite after biotite indicate retrogression from amphibolite to greenschist facies conditions. In contrast to sample TSC-56, sample TSC-13 is characterized by feldspar augen and monophase recrystallized quartz aggregates occurring in a fine-grained polymineralic matrix. The dominant recrystallization mechanism is subgrain rotation (SGR) recrystallization, possibly indicating both lower strain and lower temperatures (e.g., Stipp et al., 2002). The obtained strong Ymaxima of dynamically recrystallized quartz, however, were also found to develop at temperatures of up to ∼500◦ (e.g., Pennacchioni et al., 2010). Straight grain boundaries within the quartz aggregates point toward seizing crystal-plastic deformation at the end of the main ductile deformation under temperatures that were still high enough to enable grain boundary mobility, i.e., static grain growth (e.g., Passchier and Trouw, 2005; Herwegh et al., 2008; Platt and Behr, 2011b). Furthermore, the orthogonal orientation of stretching lineations (TSC-56 versus TSC-13, Fig. 5a) and the obtained kinematic indications (Fig. 11) cannot be interpreted as having formed under the same flow kinematics. On the basis of these considerations we regard sample TSC-56 as a representative of a high-strain zone active during top-to-the-NE extension. Chloritization of biotite might indicate retrogression due to footwall exhumation. Extension localized by partial and selective reactivation of the KTBZ and is found mainly to the west of the SR. Sample TSC-13 records a different microstructure, formed during top-to-the-NW shearing under amphibolite facies metamorphic conditions. The other two samples (TSC-28 and TSC-62) are from different lenses of (ultra-) mylonite that formed within the granitic gneiss (Fig. 5). Although they are not geographically related, they can still be regarded as more strongly deformed equivalents to TSC-13. The (ultra-) mylonites formed at the expense of the granitic gneiss, most probably under lower-grade metamorphic conditions, as indicated by protocataclastic deformation at transition zones (Fig. 10d and e). Grain size reduction in the mylonites

95

might also be an indication for lower temperature conditions and hence higher stresses due to strain localization. In contrast to previous studies (e.g., Bugge, 1928; Starmer, 1979, 1985) we interpret the mostly feldspathic augen dispersed in the fine-grained matrix (Fig. 10b and c) as clasts and not as newly grown porphyroblasts. In our interpretation they resisted intracrystalline deformation, but were dramatically rounded (Fig. 10i and d). The microstructures in samples TSC-28 show quartz ribbons with a pronounced Y-axis maximum, similar to TSC-13 (Fig. 11). In sample TSC-62, the progressive dispersion of quartz aggregates resulting in phase mixing, grain size reduction and CPO weakening goes hand in hand with a switch in deformation mechanism from dislocation creep to diffusion creep.

7. U–Pb geochronology of granite protoliths According to published geological maps (Dons and Jorde, 1987; Nordgulen, 1999) and our own observations, the c. 110 km long granitic gneiss belt hosting most of the KTBZ is exposed up to c. 5 km from the SR inside the Kongsberg lithotectonic unit and much more inside the Telemark lithotectonic unit. However, granitic gneisses have distinct fabrics on both sides of the SR. A more reddish, leucocratic fine-grained facies dominates on the Telemark side, while a coarse-grained, augen and inequigranular facies dominates on the Kongsberg side. Zircon U–Pb geochronology was performed on six samples from metamorphosed granite collected across the SR (Fig. 2b) with two specific objectives in mind: to test for (i) a possible age break at the boundary and (ii) the presence of metamorphic shear-related zircon in mylonitic samples, i.e., neoformed zircon. Samples and results are summarized from south to north and are presented in Fig. 12. BBI-09 is a specimen of the coarse-grained granitic gneiss facies along the Kolsjø section at the southern end of the Kongsberg lithotectonic unit (Fig. 2b). The sample was collected from poorly-foliated biotite-bearing granitic gneiss with centimeterscale megacrysts of a ternary K-feldspar. Zircons show a prismatic habit with generally oscillatory growth zoning (Fig. 12g). Twelve of the fourteen zircon U–Pb analyses are concordant and define a concordia age of 1158 ± 6 Ma (Fig. 12a), unequivocally recording the age of the magmatic protolith (Corfu et al., 2003). GVI-25 comes from the transition zone between granitic gneiss and mylonite in the Flesberg area (Figs. 2b, 5a). It is a red, coarse-grained augen gneiss hosting centimeter-scale rounded megacrysts of K-feldspar and plagioclase together with elongate quartz aggregates in a fine-grained matrix. Minor epidote, titanite and muscovite are present. Feldspars show protocataclastic deformation. No recrystallization or neoformation of zircon is observed, but some zircon crystals display evidence for brittle fracturing (Fig. 12h). Five analyses in prismatic crystals record crystallization of the granite protolith at 1170 ± 9 Ma (Fig. 12b). GVI-26 represents a strongly deformed granitic gneiss lens from within the banded gneiss series west of the SR (Figs. 2b, 5a). The petrography and structure of the granitic gneiss is described above by means of sample TSC-56 (Section 6.2.) Zircons show a simple oscillatory zoning with no visible evidence for reworking during deformation. Fourteen analyses in 12 zircon crystals define a concordia age at 1169 ± 5 Ma (Fig. 12c) recording crystallization of the granite precursor. BBI-02 is a red leucocratic poorly-foliated granitic gneiss collected in the Flesberg area more than 1 km west of the trace of the SR. It represents the dominant facies for the Telemark lithotectonic unit (Figs. 2b, 5a). Zircon in this sample is rich in U (generally >500 ppm) and therefore the data are scattered and several of the analyses are discordant (Fig. 12d). The six concordant analyses yield

96

T. Scheiber et al. / Precambrian Research 265 (2015) 78–103

Fig. 12. Zircon U–Pb data of six samples of 1170–1145 Ma granitoid along the KTBZ: Tera-Wasserburg concordia diagrams (a–f) and selected CL images of zircon (g–l) Samples are reported from S to N (see Fig. 2b for sample locations). Thin ellipses corresponds to SIMS analyses of zircon (2) and thick ellipses to concordia age (2). The concordia age is calculated for each sample using analyses marked by gray-shaded ellipses. Weighted average 207 Pb/206 Pb ages, calculated with the same analyses or more analyses, is reported for each sample in the figure. Position of selected SIMS analyses is indicated on CL images. Arrows on CL images point to rims and fractures (3h, i) possibly related to the deformation events recorded in the samples. Whole-rock SiO2 content, analyzed by XRF, is provided (in weight %) for most of the samples.

T. Scheiber et al. / Precambrian Research 265 (2015) 78–103

a concordia age at 1170 ± 11 Ma, recording the intrusion age of the granite protolith. GVI-23 is a strongly deformed granitic augen gneiss containing amphibole and biotite. Zircon shows magmatic oscillatory zoning. Ten analyses provide an intrusion age of 1162 ± 8 Ma for the granite protolith (Fig. 12e). A c. 10 ␮m discordant rim (too thin to be dated) is visible around some crystals (Fig. 12j–l). BBI-42 was collected in the Prestfoss area, on the Kongsberg side. It was taken in a c. 2 m thick layer of internally poorly-deformed feldspar porphyry, interlayered in a heterogeneous banded package of fine-grained gneiss with andesitic composition (Figs. 2b, 7a, 7e). The sample contains 5–10 mm automorphic ternary feldspar phenocrysts in a fine-grained matrix of quartz, feldspar, biotite and epidote. Automorphic oscillatory zoned phenocrysts of allanite (up to 2 mm long) occur. The well-preserved automorphic nature of the phenocrysts suggests a subvolcanic origin for this layer. Zircon is prismatic and oscillatory zoned (Fig. 12i). Fluid or melt inclusions parallel to the c-axis are visible in some of the zircon crystals (Corfu et al., 2003). Twelve analyses define a concordia age at 1166 ± 6 Ma, recording the magmatic crystallization of the porphyry (Fig. 12f). In synthesis, the U–Pb dates for the six samples range from 1170 ± 11 to 1158 ± 6 Ma (Fig. 12) and are identical within uncertainty. Hence the sampled granitoids can be attributed to the same magmatic event. This interval is slightly older than the published ID-TIMS zircon age of 1146 ± 5 Ma for one sample of granitic gneiss characterized by magnetite megacrysts (Bingen et al., 2003) and collected c. 5 km inside the Telemark lithotectonic unit west of Prestfoss in the same granitic gneiss belt (located in Fig. 2b). This sample probably represents a distinctly younger magmatic venue belonging to the same mappable granitic belt. Another important result of the present study is that no neoformed zircon crystals of significant size that could be attributed to deformation of the magmatic protolith were seen in the six dated samples. Only thin not-datable zircon rims (GVI-23, Fig. 12j–l) and fractures (Fig. 12h and i) were locally observed. 8. Regional-scale interpretation and discussion The integrated use of field observations, microstructural data, potential field geophysics and U–Pb zircon dates indicate a complex multiphase Sveconorwegian ductile evolution and a later brittle overprint of the KTBZ. Geochronological data provide a maximum age for deformation structures and structural data conditions of deformation and shear senses. 8.1. Implications of the geochronological data New and published zircon U–Pb geochronological results (Fig. 12; Bingen et al., 2003) indicate that the KTBZ is largely hosted within and along the eastern margin of a c. 110 km long granitic belt formed between 1170 ± 11 and 1146 ± 5 Ma. This age frame puts a limit on the onset of Sveconorwegian deformation along the boundary and, de facto, excludes pre-Sveconorwegian structural inheritance. The lack of an age difference within the KTBZ also implies that there is no juxtaposition of two far-traveled or unrelated crustal units after 1145 Ma. Bimodal “within-plate” magmatism between 1170 and 1145 Ma is voluminous and widespread in the Telemark lithotectonic unit. It includes volcanic and plutonic rocks, and felsic rocks have a typical A-type geochemical signature. Its geochemical, isotopic and geotectonic significance has been discussed abundantly in the literature (e.g., Brewer et al., 2002; Laajoki et al., 2002; Bingen et al., 2003; Andersen et al., 2007). This magmatism is also documented within the Bamble lithotectonic unit (e.g., Andersen et al., 2004a; Engvik et al., 2011), but cannot be observed farther to the east in the

97

Sveconorwegian orogen. The 1170–1145 Ma granitic belt may have intruded along a pre-existing N–S trending structure, possibly following the trend of an earlier rift-related basin structure (Fig. 13a), as documented in the Telemark lithotectonic unit (e.g., Lamminen, 2011; Spencer et al., 2014). 8.2. Sequence of deformational events We suggest that the KTBZ accommodated strain during at least four distinct ductile or ductile to brittle Sveconorwegian deformation phases, in addition to later brittle reactivation. K–Ar dating of fault gouges from within the western Kongsberg complex demonstrates that the brittle faulting can be as young as Triassic (Torgersen et al., 2015). The brittle evolution, however, is not systematically analyzed in this paper. Sketch block models with the characteristics for each deformation episode are provided in Fig. 13 and a summarizing tectonic sequence diagram is given in Fig. 14. It has to be noted that the absolute timing of the different Sveconorwegian deformational phases established in the present study remains somewhat uncertain and is based on a regional correlation with well-dated features from surrounding areas. 8.2.1. Top-to-the-W deformation Evidence for westward-directed thrusting is present within the KTBZ and has been observed in the Flesberg, Prestfoss and Sokna regions, where vergent folds and S–C fabrics at the meterscale indicate strain localization along distinct top-to-the-W shear zones under ductile conditions (Figs. 7e, 8a). Since these rocks lack evidence for earlier deformation and since they are not affected by later extensional overprint, these sites are regarded as wellpreserved examples of the earliest Sveconorwegian ductile fabric in the area, most probably formed under peak metamorphic conditions at c. 1110–1080 Ma (Arendal phase after Bingen et al., 2008a). Thrust deformation and metamorphism at this time have been reported from the Bamble lithotectonic unit (Fig. 1b; Starmer, 1991, 1993) and represent the earliest episode of high-grade metamorphism of the Sveconorwegian orogeny in southern Norway (see compilation in Bingen et al., 2008b). Elsewhere within the Kongsberg lithotectonic unit this ductile fabric is either very penetrative and corresponds to the regional foliation, or it is overprinted or even obliterated by later deformation. Where preserved, its foliation is generally steep and axial planar to tight isoclinal folds with variously N-S-plunging fold axes affecting all polymetamorphic rocks and meta-igneous and metasedimentary (cover?) units (Figs. 13b, 14). The regional structural grain within the western Kongsberg complex and the Modum complex (Fig. 2a) results mainly from this phase of westward thrust deformation (see also Starmer, 1985, 1993). The intrusive contact of the c. 1170–1145 Ma old granitic belt toward the east probably represented a primary discontinuity that might have been in part (re-) activated as a thrust (Fig. 13b), as it is known from other crustalscale thrust zones elsewhere (e.g., Johnston, 1997; Scheiber et al., 2012). However, the granitic gneiss west of the SR is also tightly folded with quartzite (Fig. 2b) and there are no indications that the entire trace of the SR corresponds to a distinct localized ductile frontal thrust of the Kongsberg onto the Telemark lithotectonic unit. In the Kolsjø section, for example, no ductile strain gradient is observed and in the Buvannet section and Prestfoss area, topto-the-W fabrics are locally strongly overprinted by later tectonic reworking. 8.2.2. Transpressional sinistral shearing Evidence for reactivation of the thrusting-related fabrics is pervasive especially within the KTBZ, particularly to the east of the brittle SR. Reactivation during a regional phase of sinistral transpression can be recognized through the presence of (1) lenses

98

T. Scheiber et al. / Precambrian Research 265 (2015) 78–103

Fig. 13. Block diagram conceptual model for the tectonic evolution of the area with emphasis on the KTBZ. Structural elements from all these deformational events are preserved along this zone. (a) Intrusion of the 1170–1145 Ma granitoids follows possibly pre-existing N-S trending structures. (b) Top-to-the-W thrusting affects the whole rock volume producing axial planar isoclinal folds. (c) Renewed deformation as a result of sinistral transpression. (d) Large-scale folding producing the Norefjell-Hønefoss fold. (e) Localized top-to-the-(N)E normal shearing and faulting. (f) Post-Sveconorwegian brittle reactivation and Oslo Rift formation. Block diagrams not to scale.

of characteristic mylonites found along the trace of the SR and W of Sokna, within and at the margin of the granitic gneiss belt (Fig. 2b). These mylonites formed at the expense of the 1170–1145 Ma granitic gneiss (Fig. 12) during sinistral strike-slip shearing at the brittle-ductile transition, as indicated by field and microstructural observations from the transition zones between granitic gneiss and mylonites (Buvannet section and Flesberg area) and shear sense indicators at various scales (Fig. 6d); (2) kinematic indicators not compatible with a top-to-the-W transport, such as sinistral

asymmetric clasts in steeply E-dipping banded gneiss in the eastern part of the Prestfoss area (Fig. 7f and g); (3) asymmetric folds affecting the gneissic main foliation which we relate genetically to transpressional sinistral shear deformation. These folds have various orientations, but fold axes plot on great circles (Figs. 4b, 5c-d) often parallel to the shear foliation in the adjacent fault rocks and thus formed as a result of shearing (e.g., Alsop and Holdsworth, 2004); (4) top-to-the-NW reverse kinematics on c. N–S trending shear planes associated with moderately to steeply SE-plunging

T. Scheiber et al. / Precambrian Research 265 (2015) 78–103

S v e c o n o r w e g i a n

Timing of Deformation Regional correlation deformation conditions

Structural expression

Affected area, regional

Compression

99

postSveconorwegian

e v o l u t i o n

Sinistral transpression

Extension

N SR

W

W

E

SW

E

NE

Oslo Rift 20 km

• • •

Penetrative ductile foliation Top-to-the-west thrusts Tight to isoclinal upright folds

•

• •

Sinistral shear foliation, particularly developed along two • N-S striking belts Top-to-the-NW thrusts (Ultra-) mylonites in granitic gneiss

• Largescale folding

•

Top-to-the-NE detachment, broadening towards the north Shear foliation in the footwall (Telemark)

•

• •

Brittle faults hosting cataclasites and gouges Quartz veins Permian dykes

Granulite- to amphibolite facies

Amphibolite- to greenschist facies

Footwall exhumation from amphibolite- to greenschist facies

Lower greenschist facies to anchizone

~1110-1080 Ma Stenian

~1020-980 Ma Stenian/Tonian

~930 Ma (Intrusion of Flå granite) - 860 Ma Tonian

Phanerozoic and younger

Main foliation within the Bamble lithotectonic unit

- Mylonite Zone, Göta Älv shear zone - Kristiansand-Porsgrunn shear zone

- Mylonite Zone - Kristiansand-Porsgrunn shear zone

Fig. 14. Tectonic sequence diagram compiling the main characteristics of deformation episodes that have affected the Kongsberg lithotectonic unit. Note that the SR represents mainly a post-Sveconorwegian phenomenon.

stretching lineations (Buvannet section, Flesberg area, Prestfoss area) formed probably also as a result of sinistral transpression. Although the mylonitic lenses contain a new shear foliation that we ascribe to this sinistral transpressive reactivation, a distinctively new foliation can hardly be recognized in adjacent gneisses. There the newly formed foliation, when developed, is generally not significantly misoriented from the regional gneissic foliation, which supports geometric reactivation of the previously formed foliation during renewed shearing (Holdsworth et al., 1997). However, the S- to SE-plunging stretching lineation and associated shear sense criteria are clearly indicative of sinistral shear. Foliation analysis and interpretation of the available potential field geophysical data in the internal part of the Kongsberg lithotectonic unit suggest that the sinistral shear zones observed in the Flesberg area continue toward the SSE, i.e., toward Saggrenda (Fig. 2). Deflection of the main foliation marks sinistral shearing on the large-scale. This sinistral shear zone corresponds to the southern part of the Saggrenda-Prestfoss mylonite zone (zone 2 in Fig. 2a) proposed by Starmer (1985). Similarly, sinistral kinematics along steep transpressional metric shear zones observed in the eastern part of the Prestfoss area is documented at many localities within the Kongsberg lithotectonic unit, especially along a zone west of the Modum complex (Fig. 2b). Thus the same deformation style can be documented for the Hokksund-Solumsmo mylonite zone (zone 4 in Fig. 2a) of Starmer (1985), which might continue into the Vardefjell shear zone (zone 5 in Fig. 2a), where sinistral kinematics has

been suggested by Bingen et al. (2001). One occurrence of mylonite west of Sokna (Figs. 2b, 8b) supports a link between these shear zones, which were severely reoriented in the area of Sokna by subsequent deformation. On the basis of these observations it can be concluded that the Kongsberg lithotectonic unit is affected by at least two major sinistral shear zones, arranged en-échelon and a number of minor similar high-strain zones that locally reactivate and obliterate the older ductile fabric (Fig. 13c), and separate the western Kongsberg complex from the Telemark lithotectonic unit to the west and the Modum complex to the east (Figs. 2b, 14). Older thrust-related fabrics are prevalent within these blocks. However, younger top-to-the-(N)W kinematics can be observed within these blocks as well, which points to a significant amount of shortening orthogonal to the strike-slip deformation zone and thus defines a transpressive regime (e.g., Dewey et al., 1998). In the area between Kongsberg and Hokksund, however, large-scale folding (Fig. 2b) indicates dextral kinematics. These structures might have formed under the same transpressional stress field, and could represent an example of strain partitioning on a large scale (e.g., Tikoff and Greene, 1997), or, alternatively, they formed at a later stage. We propose that sinistral transpression within the KTBZ and the adjoining units may be correlated with sinistral transpressional shear fabrics formed along c. N–S trending strike-slip shear belts reported from the Sveconorwegian orogen also in SW Sweden (e.g., Hageskov, 1985; Park et al., 1991; Stephens et al., 1996; Viola and Henderson, 2010). During this deformational episode a

100

T. Scheiber et al. / Precambrian Research 265 (2015) 78–103

large amount of Sveconorwegian crustal shortening would have affected the Kongsberg-, Bamble- and Idefjorden lithotectonic units, as already suggested by, e.g., Hageskov (1980). The activity of the Vardefjell shear zone (zone 5 in Fig. 2a) which might represent the NW-ward continuation of the Hokksund-Solumsmo mylonite zone (zone 4 in Fig. 2a) has been dated at 1012 ± 7 to 985 ± 5 Ma (U–Pb SIMS dating of metamorphic zircon and titanite, respectively; Bingen et al., 2008a). This age range is slightly older than the U–Pb age of metamorphic zircons (SIMS method) constraining amphibolite-facies metamorphism and deformation along the Mylonite Zone at 980 ± 13 Ma (Andersson et al., 2002) and 971 ± 8 Ma (Larson et al., 1999), and also along the Idefjordeninternal Göta Älv Shear Zone dated at 974 ± 22 Ma (Ahlin et al., 2006). The Göta Älv Shear Zone is interpreted as a sinistral strike-slip zone by Park et al. (1991). Furthermore, secondary titanite associated with top-to-the-NW deformation from the Kristiansand-Porsgrunn shear zone (Fig. 1b) has been dated at 994 ± 30 Ma (de Haas et al., 2002). This titanite also possibly formed during the same large-scale tectonic regime of sinistral transpression. However, since no direct age constraint exists for sinistral shearing affecting the Kongsberg and Modum complexes, timing remains speculative and shear deformation could have initiated already at an earlier stage, possibly during overall W-ward thrusting. 8.2.3. Large-scale folding The large-scale Norefjell-Hønefoss fold structure (Fig. 9) reorients the 1170–1145 granitic gneiss, the ductile top-to-the-W fabric as well as all structures ascribed to sinistral transpression (Fig. 13d) and thus deforms features making up the boundary zone between the Kongsberg and Telemark lithotectonic units at that time. Mylonites with preserved but reoriented sinistral shear fabric outcrop W of Sokna in the northern limb of this fold (Fig. 8b). The Norefjell-Hønefoss fold structure might have formed under the same regional stress field as the sinistral shear zones subsequent to the main sinistral shear deformation either in two distinct folding episodes or in the course of a continuous evolution, where large and relatively rigid meta-igneous bodies might have controlled its geometry (Holleia gabbro and Follum granodiorite, Fig. 9a). The fact that fold axes (domains 1–8 in Fig. 9) are distributed along a great circle even after rotation around a steeply E-plunging axis (Fig. 9c) points toward a non-cylindrical geometric evolution of this kilometric-scale fold. Evidence for a continuation of this fold structure along strike in the Idefjorden lithotectonic unit east of the Oslo Rift is not straightforward, when looking at the foliation trace map of Hageskov (1980). Parts of the fold structure are probably hidden by Paleozoic sedimentary rocks inside the Oslo Rift (Fig. 14), because the graben boundary is not a vertical structure but there is an onlap of Paleozoic rocks onto the Sveconorwegian basement (Ebbing et al., 2007). The existence of this fold strongly supports the idea that rocks of the Kongsberg lithotectonic unit continue toward the NW, north of the Vardefjell shear zone (zone 5 in Fig. 2a), as previously suggested by Nordgulen (1999) and Andersen et al. (2002). Unfortunately, the contact to the Idefjorden lithotectonic unit toward the NE (Idefjorden lithotectonic unit east of the Oslo Rift) is not clear and its nature has to be addressed by further studies. 8.2.4. Top-to-the-E/NE extension In the Buvannet, Flesberg and Prestfoss areas the immediate west of the SR is characterized by a generally E-dipping gneissic foliation with a NE-plunging stretching lineation. Associated shear sense indicators both from the field (Fig. 7d) and from the microstructural study (Fig. 11a), point to top-to-the-NE extension. Thus the trace of the SR largely coincides with an extensional shear

zone, requiring a younger exhumation of the Telemark lithotectonic unit in the footwall of this detachment (Figs. 13e, 14). The apparent eastward increase in metamorphic grade in the Telemark lithotectonic unit is a result of this exhumation. Gentle to close (N)E-vergent folds with subhorizontal fold axes and subhorizontal axial planes (Figs. 3c, 6g, 8c) are generally restricted to the west of the trace of the SR and are also related to this extensional overprint. Toward the internal part of the footwall, i.e., toward the west, the foliation progressively flattens out (Fig. 13e), and a gently SSW-plunging stretching lineation is prominent (Figs. 5 and 7). No unambiguous shear sense indicators could be assigned to this lineation. However, its orientation equals the stretching lineation formed during sinistral shear in the Kongsberg lithotectonic unit and could therefore represent remnants of the same fabric but formed at a deeper structural level. Toward the Sokna area in the north, extensional structures seem to splay and strain is accommodated within a broader zone (see Fig. 8d). This might explain why the main foliation is still preserved and not significantly obliterated, and thus the Norefjell-Hønefoss fold can still be recognized at the map-scale. In the Sokna area, normal faults cut across the older fabric at high angles (Fig. 8d). Their transport direction is toward the E to ESE, which differs from the extension direction farther south. Extension triggered the rise of pegmatitic dykes along normal faults and possibly also the ascent of the 930 Ma Flå granite intrusion (Fig. 13e), which belongs to a large-scale c. N–S trending magmatic suite, the so called Bohus-Flå Granite belt (Andersson et al., 1996). This major extensional shear zone might be interpreted as having formed in response to the late Sveconorwegian gravitational collapse (Bingen et al., 2006), as similarly reported from the Mylonite Zone (Viola et al., 2011) and from the KristiansandPorsgrunn shear zone (Mulch et al., 2005) (Fig. 1b). However, synkinematically grown micas indicate that these extensional structures were active down to 922–860 Ma and 891–880 Ma, respectively (op. cit.), and thus post-date intrusions of the BohusFlå Granite belt. 8.3. Precambrian deformation zones: typically polyphase? Due to their old age, Precambrian deformation zones (both brittle and ductile) are likely to have been reactivated multiple times during the tectonic episodes that post-date their nucleation. The mechanical weakness inherited from their earliest deformation increments thus becomes an important intrinsic property of the deforming system and it steers their subsequent geometric, kinematic and mechanic evolution (e.g., Holdsworth et al., 1997; Rutter et al., 2001; Mattila and Viola, 2014). This property plays a significant role in the process of terrane amalgamation, wherein comparatively less deformed crustal units, with probable lithospheric keel of distinct thickness, are juxtaposed along narrow and long-lived high-strain domains, which shape the first-order architecture of orogens (e.g., O’Brien et al., 1993; Subrahmanyam et al., 1993; Holdsworth and Pinheiro, 2000; Viola et al., 2008; Dufréchou et al., 2014). The KTBZ can be regarded as a prime example of an orogenscale Proterozoic deformation zone. Our combined structural and geochronological study is the first of its kind in the central part of the Sveconorwegian orogen. It adds significantly to the results of similar studies elsewhere in the orogen toward a refined understanding of the tectonic evolution (e.g., Park et al., 1991; Wahlgren et al., 1994; Stephens et al., 1996; Viola et al., 2011; Möller et al., in press). Our observations, however, also illustrate that reactivation at the large scale can be transient and very selective. It is shown that the orientation of the inherited fabrics with regard to the active stress field plays a key role in determining which volumes of the

T. Scheiber et al. / Precambrian Research 265 (2015) 78–103

deformation zone will be exploited by younger strain increments. Furthermore, temporal constraints are crucial for the reconstruction of the complete strain path of Precambrian deformation zones and whenever possible geochronology should be used. Probably the most important implication of our work is that geological studies dealing with the geometric, kinematic and temporal evolution of old deformation zones should not be based on a few, scattered observations, possibly collected only at a limited number of outcrops. They should instead be as comprehensive as possible and rely on large datasets and on a well-integrated multidisciplinary approach. 9. Conclusions Structural and geochronological data depict the KTZB as a representative example of a complexly reactivated Proterozoic deformation zone. Structural relics deriving from multiple distinct Sveconorwegian and post-Sveconorwegian deformational events overprint and crosscut each other within this zone. The Sveconorwegian structural evolution in the area can be understood in terms of four main ductile or ductile to brittle deformation phases: (i) top-to-the-W thrusting associated with regional-scale pervasive high-grade metamorphism within both the Telemark and Kongsberg lithotectonic units, (ii) sinistral transpressive shearing along two major localized shear zones, which formed mylonites characterized by pink K-feldspar augen in a fine grained matrix, (iii) large-scale complex folding affecting the earlier fabrics in both the Kongsberg and Telemark lithotectonic units in the area between Norefjell and Hønefoss, (iv) top-to-the-(N)E extensional normal shearing and faulting. Extensive post-Sveconorwegian brittle deformation localized within the KTBZ and triggered eventually the formation of the SR, the most pervasive brittle fault zone. The Norefjell-Hønefoss fold provides an elegant geometric model for the problematic “triple junction” of lithotectonic domains, invariably present on all maps, at the northern end of the Kongsberg lithotectonic unit. The presence of the fold implies that the Sveconorwegian main foliation can be traced all across from Telemark into Kongsberg and farther into the Idefjorden lithotectonic unit. The granitic gneiss on both sides of the SR has an intrusive age ranging from 1170 ± 11 to 1146 ± 5 Ma and forms a major igneous belt constraining the age of deformation along the KTBZ to be younger than 1145 Ma. In this respect, the KTBZ does not represent a terrane boundary separating two units with completely differing geologic and tectonic histories after c. 1145 Ma. This study suggests that complex reactivation of Precambrian shear zones can be anticipated to be the rule rather than the exception. Acknowledgements The Geological Survey of Norway and the counties of Telemark, Vestfold and Buskerud are warmly acknowledged for financing all mapping activities within the KONGMO project under the management of G. Viola. T. Scheiber’s contribution to the project was made possible by a post-doctoral bursary financed by the Swiss National Foundation (grant no. PBBEP2-146226) and by the support of the Geological Survey of Norway. M. Peters acknowledge financial support from the Swiss National Foundation (grant no. 200021-144381/1). SIMS data were collected at the NORDSIM laboratory, operated under an agreement between the research funding agencies of Denmark, Iceland, Norway, Sweden and Finland, the Geological Survey of Sweden and the Swedish Museum of Natural History. M. Whitehouse, L. Ilyinsky and K. Lindén guided collection

101

of SIMS data. This is NORDSIM publication #399. Iain Henderson is thanked for fruitful discussions. This article benefitted from constructive review by Michael Stephens and Saibal Gupta and helpful comments by Rüdiger Kilian.

References Adams, B.L., Wright, S.I., Kunze, K., 1993. Orientation imaging: the emergence of a new microscopy. Metall. Trans. 24A, 819–830. Ahlin, S., Hegardt, E.A., Cornell, D., 2006. Nature and stratigraphic position of the 1614 Ma Delsjön augen granite-gneiss in the Median Segment of south-west Sweden. GFF 128, 21–32. Alsop, G.I., Holdsworth, R.E., 2004. Shear zone folds: records of flow perturbation or structural inheritance? Geological Society, London, Special Publications 224, 177–199. Andersen, T., 2005. Terrane analysis, regional nomenclature and crustal evolution in the Southwest Scandinavian Domain of the Fennoscandian Shield. GFF 127, 159–168. Andersen, T., Graham, S., Sylvester, A.G., 2007. Timing and tectonic significance of Sveconorwegian A-type granitic magmatism in Telemark, southern Norway: new results from laser-ablation ICPMS U–Pb dating of zircon. NGU Bull. 447, 17–31. Andersen, T., Griffin, W.L., Jackson, S.E., Knudsen, T.L., Pearson, N.J., 2004a. MidProterozoic magmatic arc evolution at the southwest margin of the Baltic Shield. Lithos 73, 289–318. Andersen, T., Griffin, W.L., Pearson, N.J., 2002. Crustal evolution in the SW part of the Baltic shield: the Hf isotope evidence. J. Petrol. 43, 1725–1747. Andersen, T., Laajoki, K., Saeed, A., 2004b. Age, provenance and tectonostratigraphic status of the Mesoproterozoic Blefjell quartzite, Telemark sector, southern Norway. Precambrian Res. 135, 217–244. Andersson, J., Möller, C., Johansson, L., 2002. Zircon geochronology of migmatite gneisses along the Mylonite Zone (S Sweden): a major Sveconorwegian terrane boundary in the Baltic Shield. Precambrian Res. 114, 121–147. Andersson, M., Lie, J.E., Husebye, E.S., 1996. Tectonic setting of post-orogenic granites within SW Fennoscandia based on deep seismic and gravity data. Terra Nova 8, 558–566. Bezerra, F.H.R., Rossetti, D.F., Oliveira, R.G., Medeiros, W.E., Neves, B.B.B., Balsamo, F., Nogueira, F.C.C., Dantas, E.L., Andrades Filho, C., Góes, A.M., 2014. Neotectonic reactivation of shear zones and implications for faulting style and geometry in the continental margin of NE Brazil. Tectonophysics 614, 78–90. Bingen, B., Birkeland, A., Nordgulen, Ø., Sigmond, E.M.O., 2001. Correlation of supracrustal sequences and origin of terranes in the Sveconorwegian orogen of SW Scandinavia: SIMS data on zircon in clastic metasediments. Precambrian Res. 108, 293–318. Bingen, B., Davis, W.J., Hamilton, M.A., Engvik, A.K., Stein, J.J., Skår, Ø., Nordgulen, Ø., 2008a. Geochronology of high-grade metamorphism in the Sveconorwegian belt, S. Norway: U–Pb, Th–Pb and Re–Os data. Norwegian J. Geol. 88, 13–42. Bingen, B., Nordgulen, Ø., Sigmond, E.M.O., Tucker, R., Mansfeld, J., Högdahl, K., 2003. Relations between 1.19 and 1.13 Ga continental magmatism, sedimentation and metamorphism, Sveconorwegian province, S Norway. Precambrian Res. 124, 215–241. Bingen, B., Nordgulen, Ø., Viola, G., 2008b. A four-phase model for the Sveconorwegian orogeny, SW Scandinavia. Norwegian J. Geol. 88, 43–72. Bingen, B., Skår, Ø., Marker, M., Sigmond, E.M.O., Nordgulen, Ø., Ragnhildstveit, J., Mansfeld, J., Tucker, R.D., Liégeois, J.-P., 2005. Timing of continental building in the Sveconorwegian orogen, SW Scandinavia. Norwegian J. Geol. 85, 87–116. Bingen, B., Stein, H.J., Bogaerts, M., Bolle, O., Mansfeld, J., 2006. Molybdenite Re–Os dating constrains gravitational collapse of the Sveconorwegian orogen, SW Scandinavia. Lithos 87, 328–346. Bogdanova, S.V., Bingen, B., Gorbatschev, R., Kheraskova, T.N., Kozlov, V.I., Puchkov, V.N., Volozh, Y.A., 2008. The East European Craton (Baltica) before and during the assembly of Rodinia. Precambrian Res. 160, 23–45. Brewer, T.S., Åhäll, K.-I., Darbyshire, D.P.F., Menuge, J.F., 2002. Geochemistry of late Mesoproterozoic volcanism in southwestern Scandinavia: implications for Sveconorwegian/Grenvillian plate tectonic models. J. Geol. Soc. 159, 129–144. Bugge, A., 1928. En forkastning i det syd-norske grunnfjell. Norges Geologiske Undersokelse Bulletin 130, 1–124. Bugge, A., 1937. Flesberg og Eiker. NGU Bull. 143, 118 pp. Cawood, P.A., Nemchin, A.A., Strachan, R., Prave, T., Krabbendam, M., 2007. Sedimentary basin and detrital zircon record along East Laurentia and Baltica during assembly and breakup of Rodinia. J. Geol. Soc. 164, 257–275. Corfu, F., Hanchar, J.M., Hoskin, P.W.O., Kinny, P., 2003. Atlas of zircon textures. In: Hanchar, J.M., Hoskin, P.W.O. (Eds.), Reviews in Mineralogy and Geochemistry. Mineral Society of America, Washington, pp. 469–500. de Haas, G.-J., Nijland, T.G., Andersen, T., Corfu, F., 2002. New constraints on the timing of deposition and metamorphism in the Bamble sector, south Norway: zircon and titanite U–Pb data from the Nelaug area. GFF 124, 73–78. Dewey, J.F., Holdsworth, R.E., Strachan, R.A., 1998. Transpression and transtension zones. Geological Society, London, Special Publications 135, 1–14. Dons, J.A., 1960. Telemark supracrustals and associated rocks. In: Holtedahl, O. (Ed.), Geology of Norway. Norges geologiske undersøkelse, Trondheim, pp. 49–58. Dons, J.A., Jorde, K., 1987. Geologisk kart over Norge, Berggrunnskart SKIEN. NGU, Trondheim.

102

T. Scheiber et al. / Precambrian Research 265 (2015) 78–103

Dufréchou, G., Harris, L.B., Corriveau, L., 2014. Tectonic reactivation of transverse basement structures in the Grenville orogen of SW Quebec, Canada: insights from gravity and aeromagnetic data. Precambrian Res. 241, 61–84. Ebbing, J., Skilbrei, J.R., Olesen, O., 2007. Insights into the magmatic architecture of the Oslo Graben by petrophysically constrained analysis of the gravity and magnetic field. J. Geophys. Res. Solid Earth 112, B04404. Engvik, A.K., Mezger, K., Wortelkamp, S., Bast, R., Corfu, F., Korneliussen, A., Ihlen, P., Bingen, B., Austrheim, H., 2011. Metasomatism of gabbro – mineral replacement and element mobilization during the Sveconorwegian metamorphic event. J. Metamorph. Geol. 29, 399–423. Govers, R., Wortel, M.J.R., 1995. Extension of stable continental lithosphere and the initiation of lithospheric scale faults. Tectonics 14, 1041–1055. Hageskov, B., 1980. The Sveconorwegian structures of the Norwegian part of the Kongsberg-Bamble-Østfold segment. GFF 102, 150–155. Hageskov, B., 1985. Constructional deformation of the Koster dyke swarm in a ductile sinistral shear zone, Koster island, SW Sweden. Bull. Geol. Soc. Denmark 34, 151–197. Heilbronner, R., Tullis, J., 2006. Evolution of c axis pole figures and grain size during dynamic recrystallization: results from experimentally sheared quartzite. J. Geophys. Res. Solid Earth 111, B10202. Herwegh, M., Berger, A., Ebert, A., Brodhag, S., 2008. Discrimination of annealed and dynamic fabrics: consequences for strain localization and deformation episodes of large-scale shear zones. Earth Planet. Sci. Lett. 276, 52–61. Hippertt, J., Rocha, A., Lana, C., Egydio-Silva, M., Takeshita, T., 2001. Quartz plastic segregation and ribbon development in high-grade striped gneisses. J. Struct. Geol. 23, 67–80. Holdsworth, R.E., Butler, C.A., Roberts, A.M., 1997. The recognition of reactivation during continental deformation. J. Geol. Soc. 154, 73–78. Holdsworth, R.E., Pinheiro, R.V.L., 2000. The anatomy of shallow-crustal transpressional structures: insights from the Archaean Carajás fault zone, Amazon, Brazil. J. Struct. Geol. 22, 1105–1123. Holdsworth, R.E., Stewart, M., Imber, J., Strachan, R.A., 2001. The structure and rheological evolution of reactivated continental fault zones: a review and case study. Geological Society, London, Special Publications 184, 115–137. Jacobsen, S.B., Heier, K.S., 1978. Rb-Sr isotope systematics in metamorphic rocks, Kongsberg sector, south Norway. Lithos 11, 257–276. Jakobsen, P.R., 1991. Sokna-Områdets Grundfjeldsgeologi. Københavns Universitet, pp. 54. Johnston, J.D., 1997. Localization of mid-crustal thrust ramps by metadolerite sheets in the Dalradian of northwest Ireland. Geol. Mag. 134, 199–212. Kilian, R., Heilbronner, R., Stünitz, H., 2011. Quartz grain size reduction in a granitoid rock and the transition from dislocation to diffusion creep. J. Struct. Geol. 33, 1265–1284. Köykkä, J., Lamminen, J., 2011. Tidally influenced clastic epeiric sea at a Mesoproterozoic continental margin, Rjukan Rift Basin, southern Norway. Precambrian Res. 185, 164–182. Laajoki, K., Corfu, F., Andersen, T., 2002. Lithostratigraphy and U–Pb geochronology of the Telemark supracrustals in the Bandak-Sauland area, Telemark, South Norway. Norwegian J. Geol. 2002, 119–138. Lamminen, J., 2011. Provenance and correlation of sediments in Telemark, South Norway: status of the Lifjell Group and implications for early Sveconorwegian fault tectonics. Norwegian J. Geol. 91, 57–75. Lamminen, J., Köykkä, J., 2010. The provenance and evolution of the Rjukan Rift Basin, Telemark, south Norway: the shift from a rift basin to an epicontinental sea along a Mesoproterozoic supercontinent. Precambrian Res. 181, 129–149. Larson, S.Å., Cornell, D.H., Armstrong, R.A., 1999. Emplacement ages and metamorphic overprinting of granitoids in the Sveconorwegian Province in Varmland, Sweden - an ion probe study. Nord. Geol. Tidsskr. 79, 87–95. Law, R.D., 1990. Crystallographic fabrics: a selective review of their applications to research in structural geology. Geological Society, London, Special Publications 54, 335–352. Li, Z.X., Bogdanova, S.V., Collins, A.S., Davidson, A., De Waele, B., Ernst, R.E., Fitzsimons, I.C.W., Fuck, R.A., Gladkochub, D.P., Jacobs, J., Karlstrom, K.E., Lu, S., Natapov, L.M., Pease, V., Pisarevsky, S.A., Thrane, K., Vernikovsky, V., 2008. Assembly, configuration, and break-up history of Rodinia: a synthesis. Precambrian Res. 160, 179–210. Ludwig, K.R., 1998. On the treatment of concordant uranium-lead ages. Geochim. Cosmochim. Acta 62, 665–676. Ludwig, K.R., 2008. User’s Manual for Isoplot 3.70. A Geochronological Toolkit for Microsoft Excel. Berkley Geochronology Center, Special Publication 4. Madritsch, H., Kounov, A., Schmid, S.M., Fabbri, O., 2009. Multiple fault reactivations within the intra-continental Rhine–Bresse Transfer Zone (La Serre Horst, eastern France). Tectonophysics 471, 297–318. Mancktelow, N.S., Pennacchioni, G., 2005. The control of precursor brittle fracture and fluid–rock interaction on the development of single and paired ductile shear zones. J. Struct. Geol. 27, 645–661. Mattila, J., Viola, G., 2014. New constraints on 1.7 Gyr of brittle tectonic evolution in southwestern Finland derived from a structural study at the site of a potential nuclear waste repository (Olkiluoto Island). J. Struct. Geol. 67A, 50–74. Mirabella, F., Ciaccio, M.G., Barchi, M.R., Merlini, S., 2004. The Gubbio normal fault (Central Italy): geometry, displacement distribution and tectonic evolution. J. Struct. Geol. 26, 2233–2249. Mulch, A., Cosca, M.A., Andresen, A., Fiebig, J., 2005. Time scales of deformation and exhumation in extensional detachment systems determined by highspatial resolution in situ UV-laser 40Ar/39Ar dating. Earth Planet. Sci. Lett. 233, 375–390.

Munz, I.A., Wayne, D., Austrheim, H., 1994. Retrograde fluid infiltration in the high-grade Modum Complex South Norway: evidence for age, source and REE mobility. Contrib. Mineral. Petrol. 116, 32–46. Möller, C., Andersson, J., Dyck, B., Lundin, I.A. Exhumation of an eclogite terrane as a hot migmatitic nappe, Sveconorwegian orogen. Lithos (in press). Möller, C., Bingen, B., Andersson, J., Stephens, M.B., Viola, G., Scherstén, A., 2013. A non-collisional, accretionary Sveconorwegian orogen – Comment. Terra Nova 25, 165–168. Nordgulen, Ø., 1999. Geologisk kart over Norge, Berggrunnskart HAMAR. NGU, Trondheim. O’Brien, B.H., O’Brien, S.J., Dunning, G.R., Tucker, R.D., 1993. Episodic reactivation of a Late Precambrian mylonite zone on the Gondwanan Margin of the Appalachians, southern Newfoundland. Tectonics 12, 1043–1055. Occhipinti, S.A., Reddy, S.M., 2004. Deformation in a complex crustal-scale shear zone: Errabiddy Shear Zone, Western Australia. Geological Society, London, Special Publications 224, 229–248. Park, R.G., Åhäll, K.I., Bland, M.P., 1991. The Sveconorwegian shear-zone network of SW Sweden in relation to mid-Proterozoic plate movements. Precambrian Res. 49, 245–260. Passchier, C.W., Hoek, J.D., Bekendam, R.F., De Boorder, H., 1990. Ductile reactivation of Proterozoic brittle fault rocks; an example from the Vestfold Hills, East Antarctica. Precambrian Res. 47, 3–16. Passchier, C.W., Trouw, R.A.J., 2005. Microtectonics. Springer Verlag, Berlin. Paterson, M.S., 1995. A theory for granular flow accommodated by material transfer via an intergranular fluid. Tectonophysics 245, 135–151. Pedersen, S., Andersen, T., Konnerup-Madsen, J., Griffin, W., 2009. Recurrent mesoproterozoic continental magmatism in South-Central Norway. Int. J. Earth. Sci. 98, 1151–1171. Pennacchioni, G., Menegon, L., Leiss, B., Nestola, F., Bromiley, G., 2010. Development of crystallographic preferred orientation and microstructure during plastic deformation of natural coarse-grained quartz veins. J. Geophys. Res. Solid Earth 115, B12405. Platt, J.P., Behr, W.M., 2011a. Deep structure of lithospheric fault zones. Geophys. Res. Lett. 38, L24308. Platt, J.P., Behr, W.M., 2011b. Grain size evolution in ductile shear zones: implications for strain localization and the strength of the lithosphere. J. Struct. Geol. 33, 537–550. Prucha, J.J., 1992. Zone of weakness concept: a review and evaluation. In: Bartholomew, M.J., Hyndman, D.W., Mogk, D.W., Mason, R. (Eds.), Basement Tectonics 8. Springer, The Netherlands, pp. 83–92. Roberts, N.M.W., Slagstad, T., 2014. Continental growth and reworking on the edge of the Columbia and Rodinia supercontinents: 1.86–0.9 Ga accretionary orogeny in southwest Fennoscandia. Int. Geol. Rev., 1–25. Rohr-Torp, E., 1973. Permian rocks and faulting in Sandsvær at the western margin of the Oslo region. NGU Bull. 300, 53–71. Romer, R.L., 1996. Contiguous Laurentia and Baltica before the GrenvillianSveconorwegian orogeny? Terra Nova 8, 173–181. Rutter, E.H., Holdsworth, R.E., Knipe, R.J., 2001. The nature and tectonic significance of fault-zone weakening: an introduction. Geological Society, London, Special Publications 186, 1–11. Røhr, T.S., Bingen, B., Robinson, P., Reddy, S.M., 2013. Geochronology of Paleoproterozoic augen gneisses in the Western Gneiss region, Norway: evidence for Sveconorwegian zircon neocrystallization and Caledonian zircon deformation. J. Geol. 121, 105–128. Scheiber, T., Pfiffner, O.A., Schreurs, G., 2012. Strain accumulation during basal accretion in continental collision – a case study from the Suretta nappe (eastern Swiss Alps). Tectonophysics 579, 56–73. Schmid, S.M., Casey, M., 1986. Complete fabric analysis of some commonly observed quartz C-axis patterns. In: Mineral and Rock Deformation: Laboratory Studies. AGU, Washington, DC, pp. 263–286. Scholz, C.H., 1988. The brittle-plastic transition and the depth of seismic faulting. Geol. Rundsch. 77, 319–328. Sibson, R.H., 1977. Fault rocks and fault mechanisms. J. Geol. Soc. 133, 191–213. Sibson, R.H., 1986. Earthquakes and rock deformation in crustal fault zones. Annu. Rev. Earth Planet. Sci. 14, 149–175. Sigmond, E.M.O., 1985. The Mandal-Ustaoset line, a newly discovered major fault zone in south Norway. In: Tobi, A.C., Touret, J.L. (Eds.), The Deep Proterozoic Crust in the North Atlantic Provinces. Reidel, Dordrecht, pp. 323–331. Sigmond, E.M.O., 1998. Geologisk kart over Norge, Berggrunnskart ODDA. NGU, Trondheim. Simpson, C., Whitmeyer, S.J., De Paor, D.G., Gromet, L.P., Miro, R., Krol, M.A., Short, H., 2001. Sequential ductile to brittle reactivation of major fault zones along the accretionary margin of Gondwana in Central Argentina. Geological Society, London, Special Publications 186, 233–255. Slagstad, T., Roberts, N.M.W., Marker, M., Røhr, T.S., Schiellerup, H., 2013. A noncollisional, accretionary Sveconorwegian orogen. Terra Nova 25, 30–37. Spencer, C.J., Roberts, N.M.W., Cawood, P.A., Hawkesworth, C.J., Prave, A.R., Antonini, A.S.M., Horstwood, M.S.A., 2014. Intermontane basins and bimodal volcanism at the onset of the Sveconorwegian Orogeny, southern Norway. Precambrian Res. 252, 107–118. Starmer, I.C., 1979. The Kongsberg Series Margin and Its Major Bend in the Flesberg Area, Numedal, Buskerud. NGU Bull. 351, 99–120. Starmer, I.C., 1985. The evolution of the South Norwegian Proterozoic as revealed by the major and mega-tectonics of the Kongsberg and Bamble sectors. In: Tobi, A.C., Touret, J.L. (Eds.), The Deep Proterozoic Crust in the North Atlantic Provinces. Reidel, Dordrecht, pp. 259–290.

T. Scheiber et al. / Precambrian Research 265 (2015) 78–103 Starmer, I.C., 1991. The Proterozoic evolution of the Bamble Sector shear belt, southern Norway: correlations across southern Scandinavia and the Grenvillian controversy. Precambrian Res. 49, 107–139. Starmer, I.C., 1993. The Sveconorwegian orogeny in Southern Norway, relative to deep-crustal structures and events in the North-Atlantic Proterozoic supercontinent. Nord. Geol. Tidsskr. 73, 109–132. Stephens, M.B., Wahlgren, C.-H., Weijermars, R., Cruden, A.R., 1996. Left-lateral transpressive deformation and its tectonic implications, Sveconorwegian orogen, Baltic Shield, southwestern Sweden. Precambrian Res. 79, 261–279. Stipp, M., Stünitz, H., Heilbronner, R., Schmid, S.M., 2002. The eastern Tonale fault zone: a ‘natural laboratory’ for crystal plastic deformation of quartz over a temperature range from 250 to 700 ◦ C. J. Struct. Geol. 24, 1861–1884. Stünitz, H., Gerald, J.D.F., 1993. Deformation of granitoids at low metamorphic grade: II. Granular flow in albite-rich mylonites. Tectonophysics 221, 299–324. Subrahmanyam, V., Ramana, M.V., Rao, D.G., 1993. Reactivation of Precambrian faults on the southwestern continental margin of India: evidence from gravity anomalies. Tectonophysics 219, 327–339. Tikoff, B., Greene, D., 1997. Stretching lineations in transpressional shear zones: an example from the Sierra Nevada Batholith, California. J. Struct. Geol. 19, 29–39. Torgersen, E., Viola, G., Zwingmann, H., Henderson, I.H.C., 2015. Inclined K–Ar illite age spectra in brittle fault gouges: effects of fault reactivation and wall-rock contamination. Terra Nova. Toy, V.G., Prior, D.J., Norris, R.J., 2008. Quartz fabrics in the Alpine Fault mylonites: influence of pre-existing preferred orientations on fabric development during progressive uplift. J. Struct. Geol. 30, 602–621. Viola, G., Henderson, I.C., 2010. Inclined transpression at the toe of an arcuate thrust: an example from the Precambrian ‘Mylonite Zone’ of the Sveconorwegian orogen. Geological Society, London, Special Publications 335, 715–737.

103

Viola, G., Henderson, I.H.C., Bingen, B., Hendriks, B.W.H., 2011. The Grenvillian–Sveconorwegian orogeny in Fennoscandia: back-thrusting and extensional shearing along the Mylonite Zone. Precambrian Res. 189, 368–388. Viola, G., Henderson, I.H.C., Bingen, B., Thomas, R.J., Smethurst, M.A., de Azavedo, S., 2008. Growth and collapse of a deeply eroded orogen: insights from structural, geophysical, and geochronological constraints on the Pan-African evolution of NE Mozambique. Tectonics 27, TC5009. Viola, G., Kounov, A., Andreoli, M.A.G., Mattila, J., 2012. Brittle tectonic evolution along the western margin of South Africa: more than 500 Myr of continued reactivation. Tectonophysics 514–517, 93–114. Wahlgren, C.-H., Cruden, A.R., Stephens, M.B., 1994. Kinematics of a major fan-like structure in the eastern part of the Sveconorwegian orogen, Baltic Shield, southcentral Sweden. Precambrian Res. 70, 67–91. Whitehouse, M.J., Kamber, B.S., 2005. Assigning dates to thin gneissic veins in high-grade metamorphic terranes: a cautionary tale from Akilia, Southwest Greenland. J. Petrol. 46, 291–318. Whitehouse, M.J., Kamber, B.S., Moorbath, S., 1999. Age significance of U–Th–Pb zircon data from early Archaean rocks of west Greenland – a reassessment based on combined ion-microprobe and imaging studies. Chem. Geol. 160, 201–224. Wiedenbeck, M., AllÉ., P., Corfu, F., Griffin, W.L., Meier, M., Oberli, F., Quadt, A.V., Roddick, J.C., Spiegel, W., 1995. Three natural zircon standards for U–Th–Pb, Lu–Hf, trace element and REE analyses. Geostand. Newsl. 19, 1–23. Ziegler, P.A., 1996. Geodynamic processes governing development of rifted basins. In: Roure, F., Ellouz, N., Shein, V.S., Skvortsov, I. (Eds.), Geodynamic Evolution of Sedimentary Basins. Inst. Fr. du Pét./Technip, Paris, pp. 19–67.