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# The Sveconorwegian orogeny

Bernard Bingen, Giulio Viola, Charlotte Möller, Jacqueline Vander Auwera, Antonin Laurent, Keewook Yi

► To cite this version:

Bernard Bingen, Giulio Viola, Charlotte Möller, Jacqueline Vander Auwera, Antonin Laurent, et al. The Sveconorwegian orogeny. Gondwana Research, 2021, 90, pp.273-313. 10.1016/j.gr.2020.10.014. hal-04922898

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Distributed under a Creative Commons Attribution 4.0 International License
2 The Sveconorwegian orogeny

3

4

5 Bernard Bingen¹, Giulio Viola², Charlotte Möller³, Jacqueline Vander Auwera⁴, Antonin

6 Laurent⁵, Keewook Yi⁶

7

8

9 1 Corresponding author, Geological Survey of Norway, 7491 Trondheim, Norway;
10 bernard.bingen@ngu.no; ORCID; 0000-0002-4901-2016

11 2 Department of Biological, Geological and Environmental Sciences, BiGeA, University of
12 Bologna, 40126 Bologna, Italy; giulio.viola3@unibo.it

13 3 Department of Geology, Lund University, Sölvegatan 12, SE-223 62 Lund, Sweden;
14 charlotte.moller@geol.lu.se

15 4 Department of Geology, University of Liège, 4000 Liège, Belgium; jvdauwera@uliege.be

16 5 Université de Lyon, UJM-Saint-Etienne, CNRS, UCA, IRD, LMV UMR 6524, F-42023
17 Saint-Etienne, France; antonin.laurent@univ-st-etienne.fr

18 6 Korea Basic Science Institute, 363-883 Chungbuk, South Korea; kyi@kbsi.re.kr

19

20

21 **GRAPHICAL ABSTRACT**
Sveconorwegian orogeny

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ABSTRACT

This article reviews the geology of the Sveconorwegian orogen in south Scandinavia and existing tectonic models for the Mesoproterozoic to Neoproterozoic Sveconorwegian orogeny. It proposes an updated geodynamic scenario of large, hot, long-duration continental collision starting at c. 1065 Ma between proto-Baltica and another plate, presumably Amazonia, in a Rodinia-forming context. An orogenic plateau formed at 1280 Ma as a back-arc Cordillera-style plateau, and then grew further stepwise after 1065 Ma, as a collisional Tibetan-style plateau. Voluminous mantle- and crustal-derived Sveconorwegian magmatism took place in the hinterland in the west of the orogen, mainly: (i) bimodal magmatism at 1280–1145 Ma, overlapping with extensional intramontane basin sedimentation, (ii) the calc-alkaline Sirdal magmatic belt at 1065–1020 Ma, (iii) the hydrous ferroan hornblende-biotite granite (HBG) suite at 985–925 Ma and (iv) the anhydrous ferroan massif-type anorthosite-mangerite-charnockite (AMC) suite at 935–915 Ma. High-alumina orthopyroxene megacrysts
36 in anorthosite imply mafic underplating at 1040 Ma and remelting of the underplates at 930
37 Ma. Overlapping with magmatism, protracted low-pressure, granulite-facies metamorphism
38 reached twice ultra-high temperature conditions, of 0.6 GPa-920 °C at 1030-1005 Ma and 0.4
39 GPa-920 °C at 930 Ma. These features imply shallow asthenosphere under the crust. Towards
40 the foreland in the east, metamorphism shows increasing high-pressure signature eastwards
41 with time, with peak P-T values of 1.15 GPa-850 °C at 1150-1120 Ma in the Bamble-
42 Kongsberg lithotectonic units, 1.5 GPa-740 °C at c. 1050 Ma in the Idefjorden lithotectonic
43 unit, and 1.8 GPa-870 °C at c. 990 Ma in the Eastern Segment under eclogite-facies
44 conditions. These are attributed to retreating delamination of the dense sub-continental
45 lithospheric mantle and growth of the orogenic plateau towards the foreland. After c. 930 Ma,
46 convergence came to a halt, the orogenic plateau collapsed, and 16 km of overburden was
47 removed by extension and erosion.

48
49 **Keywords:** Sveconorwegian, Mesoproterozoic, Rodinia, continental collision, orogenic
50 plateau - lithospheric mantle delamination

51
52 **Highlights:**
53

- Review of the geology of the Sveconorwegian orogen in south Scandinavia.

54 - Review of geodynamic models for the Mesoproterozoic to Neoproterozoic Sveconorwegian orogeny.

55 - Model of large, hot, long-duration continental collision for the Sveconorwegian orogeny.

57 - Orogenic plateau construction is associated with retreating delamination of the continental lithospheric mantle.

59 - Protracted shallow asthenosphere lead to crustal melting and ultra-high temperature
60 granulite-facies metamorphism.
61 - Massif-type anorthosite-mangerite-charnockite plutonism resulted from remelting of mafic
62 underplates at 1.1 GPa under high heat flow conditions.

63 - The Sveconorwegian orogeny contributed to assembly of Rodinia supercontinent

64 **1 Introduction**

65 Late-Mesoproterozoic orogenic belts are interpreted as products of the closure of oceanic
66 realms and the collision between continents to form supercontinent Rodinia at the end of the
67 Mesoproterozoic (Hoffman, 1991; Li et al., 2008). The Rodinia paradigm is robust, and
68 supported by a peak in the abundance of late Mesoproterozoic detrital zircons (Hawkesworth
69 et al., 2009). This notwithstanding, paleogeographic models for Rodinia configuration and
70 plate tectonic models for Rodinia assembly remain in essence ill-defined (Torsvik, 2003).

71 Proto-Baltica (Proterozoic Baltica = East European Craton, here after called Baltica) is a core
72 piece of Rodinia in almost all models (Fig. 1) (Li et al., 2008; Merdith et al., 2017), and the
73 Sveconorwegian orogen at the western margin of Baltica provides key geological evidence for
74 the assembly of Rodinia (Bingen et al., 2008a; Bingen et al., 2008c; Bogdanova et al., 2008;
75 Cawood and Pisarevsky, 2017; Cawood et al., 2010; Falkum and Petersen, 1980; Gee et al.,
76 2015; Gower et al., 2008; Hartz and Torsvik, 2002; He et al., 2018; Ibanez-Mejia et al., 2011;
77 Lorenz et al., 2012; Pisarevsky et al., 2014; Roberts, 2013; Roberts and Slagstad, 2015;
78 Slagstad et al., 2019; Slagstad et al., 2020; Slagstad et al., 2018; Slagstad et al., 2017;
79 Stephens and Wahlgren, 2020b; Torsvik et al., 1996; Weber et al., 2010).

80 The Sveconorwegian orogen is well exposed and accessible in its type area in southwest
81 Scandinavia (south Norway and southwest Sweden). It represents therefore an excellent
82 natural laboratory to study Precambrian geodynamics (Bingen and Viola, 2018; Laurent et al.,
83 2018a; Möller and Andersson, 2018; Slagstad et al., 2018; Stephens and Wahlgren, 2020b;
84 Vander Auwera et al., 2011; Viola and Henderson, 2010).
The body of geological data on the Sveconorwegian orogen has been steadily growing over the last 20 years, leading to contrasting conceptual models. This article reviews the existing structural, metamorphic, magmatic, geochronological and isotopic record across the entire Sveconorwegian orogen, and discusses the orogenic models that have been proposed in the literature. In fact, there is a lively debate in the literature on whether the Sveconorwegian orogeny was a collisional or a non-collisional (Andean) orogeny (Bingen et al., 2008a; Möller and Andersson, 2018; Slagstad et al., 2020; Slagstad et al., 2017; Slagstad et al., 2013; Stephens and Wahlgren, 2020b). We address this debate and conclude proposing an updated model of large, hot, and long-duration continent-continent collision for the Sveconorwegian orogeny at the margin of Baltica / Fennoscandia. This model involves the stepwise propagation of an orogenic plateau towards the foreland and hinterland of the orogen, associated with retreating delamination of the continental lithospheric mantle. It takes into account a number of key features of the orogeny, including the zoning of metamorphism, the distribution of magmatism and the genesis of massif-type anorthosites.

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**Figure 1.** Archetypal paleogeographic reconstruction of proto-Baltica (Baltica), Laurentia and Amazonia in their Rodinia framework at the Mesoproterozoic-Neoproterozoic boundary (Cawood and Pisarevsky, 2017; Hoffman, 1991; Li et al., 2008). The first order architecture of the Meso- to Neoproterozoic orogenic belts is shown, with emphasis on the geochronology of metamorphism (Hynes and Rivers, 2010; Ibanez-Mejia et al., 2011; Rivers, 2008; Tohver et al., 2005). The high-pressure (HP) metamorphic belts are shown separately. The names of the main tectonometamorphic phases in the different orogens are listed in the legend, with Arendal, Agder, Falkenberg for the Sveconorwegian orogen, Shawinigan, Ottawan, Rigolet for the Grenville orogen, Putumayo for the Putumayo orogenic belt, Zapotecan for the Oaxaquia lithotectonic unit, and Sunsás for the Sunsás, Aguapei and N Brasilândia belts.

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**Table 1.** Chart of geological events in the Sveconorwegian orogen.

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**Figure 2.** Sketch map of the Sveconorwegian orogen, with nomenclature of lithotectonic units and main shear and fault zones.

## 2 Context

### 2.1 The Sveconorwegian orogen and Sveconorwegian orogeny

The Sveconorwegian orogen is located along the southwestern margin of Fennoscandia, which is the northern part of proto-Baltica (Fig. 1; Fig. 2; Table 1) (Bogdanova et al., 2008; Koistinen et al., 2001; Stephens et al., 2020). The Sveconorwegian orogen consists of Paleoproterozoic to Mesoproterozoic continental lithosphere reworked during the Sveconorwegian orogeny at the transition between the Mesoproterozoic and the Neoproterozoic (Stenian to Tonian). This lithosphere was generated during the Svecokarelian (1910–1750 Ma), post-Svecokarelian (1710–1660 Ma), Gothian (1660–1520 Ma), Telemarkian (1520–1480 Ma) and Hallandian (1465–1380 Ma) accretionary orogenies.
The exposed Sveconorwegian orogen is presently c. 550 km wide and has a general N-S structural grain (Fig. 2) (Berthelsen, 1980; Demaiffe and Michot, 1985; Falkum, 1985; Falkum and Petersen, 1980). In the east, it is separated from the Paleoproterozoic foreland by the nearly 700 km long Sveconorwegian front (Möller and Andersson, 2018; Möller et al., 2015; Stephens and Wahlgren, 2020a; Wahlgren et al., 1994).

In the north, the Sveconorwegian orogen was reworked during the Caledonian orogeny (Fig. 2). Precambrian rocks with a Meso- to Neoproterozoic overprint are observed in the Western Gneiss Region, the largest basement window in the Caledonides (Røhr et al., 2013; Tucker et al., 1990) and are also found in Caledonian thrust nappes of the Lower and Middle Allochthons of the Caledonides (Augland et al., 2014; Corfu, 2019; Lundmark and Corfu, 2008; Roffeis and Corfu, 2014; Wiest et al., 2018). In the south, the Sveconorwegian basement is overlain by Phanerozoic sedimentary rocks and affected by Carboniferous-Permian and younger faulting and rifting along the WNW-ESE trending Sorgenfrei-Tornquist Fault Zone and NNE-SSW trending Oslo rift (Fig. 2) (Bergerat et al., 2007; Erlström, 2020; Larsen et al., 2008; Torgersen et al., 2015). As inferred from geophysical data and a few deep wells in Denmark, a Sveconorwegian basement probably underlies the Norwegian-Danish Basin (Ringkøbing-Fyn high), reaching the southern boundary of the Baltica plate (Trans-European Suture Zone and Elbe line) (Lassen and Thybo, 2012; Olesen et al., 2004; Olivarius et al., 2015; Thybo, 2001).

The Sveconorwegian orogen can be conceptually subdivided into five, orogen-parallel lithotectonic units (INSPIRE_Directive, 2007), called, from east to west, the Eastern Segment, and the Idefjorden, Kongsberg, Bamble, and Telemarkia lithotectonic units (also referred to as units in short in the following text) (Fig. 2) (Bingen et al., 2008c). These lithotectonic units are separated by major Sveconorwegian shear zones and are characterized by distinct geological histories.
The first high-grade metamorphism attributed to the Sveconorwegian orogeny dates back to between 1150 and 1120 Ma and is recorded in the Bamble and Kongsberg lithotectonic units. It is referred to as the Arendal phase in Bingen et al. (2008a; 2008c). As elaborated further below, this early-Sveconorwegian event can be interpreted as the outcome of a geodynamic evolution starting after the Hallandian orogeny, i.e. after c. 1340 Ma, and hereafter referred to as the pre-Sveconorwegian. The main Sveconorwegian orogeny started at c. 1065 Ma, and can be summarized by three orogenic phases (Bingen et al., 2008a; Bingen et al., 2008c): the Agder phase (1065–1000 Ma), the Falkenberg phase (1000–970 Ma) and the Dalane phase (970–900 Ma). As more geological data become available, however, these three phases are becoming increasingly difficult to discriminate in time and they are not used systematically in the following text. Intrusion of pegmatite fields and lamprophyre dykes sealed the orogeny at c. 915–900 Ma (Müller et al., 2015; Müller et al., 2017; Wahlgren et al., 2015).

## 2.2 Rodinia assembly

Several paleogeographic and tectonic models have been proposed for the configuration and assembly of supercontinent Rodinia at the end of the Mesoproterozoic (Hoffman, 1991; Li et al., 2008; Merdith et al., 2017; Torsvik, 2003). Classical models (Fig. 1), integrating paleomagnetic data and geological information from the Proterozoic to the Phanerozoic, suggest that Rodinia formed by the reassembly of continents previously assembled into supercontinent Nuna (Columbia) during the Paleo- and Mesoproterozoic (Evans and Mitchell, 2011; Johansson, 2009; Pisarevsky et al., 2014; Rogers and Santosh, 2002; Zhang et al., 2012). These models locate Laurentia in the centre of Rodinia, with Baltica to the east and Amazonia south of Laurentia, respectively (Fig. 1) (Cawood and Pisarevsky, 2017; Dalziel, 1997; Gong et al., 2018; Hoffman, 1991; Li et al., 2008; Merdith et al., 2017; Torsvik et al.,
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1996). Alternative Baltica–Laurentia reconstructions are proposed by Torsvik (2003), Lorenz et al. (2012) and Slagstad et al. (2019).

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It is beyond the scope of this paper to review Rodinia assembly models. In the following text and in several figures, updated geological, geochronological and isotopic data from the Sveconorwegian orogen are compared with data from the Grenville orogen of Laurentia, the Putumayo and Sunsás orogens of Amazonia and Mesoproterozoic lithotectonic units in the Andes (Garzón, Las Minas inliers) and Mexico (Oaxaquia). The goal is to show that it is realistic to consider the Sveconorwegian orogen as part of a large orogenic zone between Laurentia, Amazonia and Baltica (Fig. 1). The comparative analysis offers a broader perspective for the Sveconorwegian orogeny in a Rodinia context.

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**Figure 3.** Conceptual tectonic models, in map view, for the Sveconorwegian orogeny, reviewed in this paper. Same colour coding as in Fig. 2. (a) Early-Sveconorwegian accretion of the Telemarkia lithotectonic unit, with suturing along the Bamble-Kongsberg lithotectonic units (Bingen et al., 2005). (b) Wrench tectonics involving large-strike-slip displacements

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193 between the five lithotectonic units of the orogen, all endemic to Fennoscandia (Lamminen and Köykkä, 2010; Stephens and Wahlgren, 2020b). (c) Collision at c. 990 Ma between 194 Baltica (Fennoscandia) and a continent comprising the four western lithotectonic units of the 195 orogen (Möller and Andersson, 2018; Petersson et al., 2015b). (d) Collision at c. 1065 Ma 196 between Baltica (Fennoscandia) and another continental plate (Amazonia) with closure of 197 oceanic basins to the west of the exposed orogen (Bogdanova et al., 2008; Cawood and 198 Pisarevsky, 2017; Ibanez-Mejia et al., 2011). (e, f) non-collisional (Andean) models, with 199 orogeny controlled by a subduction system outboard of Fennoscandia during the entire 200 duration of orogeny from c. 1150 to 900 Ma. (e) The margin was either well assembled before 201 the orogeny (Falkum and Petersen, 1980; Slagstad et al., 2013) or (f) dispersed and re- 202 assembled during the orogeny (Slagstad et al., 2020). 203

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## 2.3 A diversity of orogenic models

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206 Many large-scale tectonic models have been proposed to explain the Sveconorwegian 207 orogenic evolution. In Fig. 3, six possible conceptual end-member models are sketched in 208 map view. They range from collisional (Himalaya-Tibet type) to non-collisional (Andean 209 type), and some involve accretion of exotic lithotectonic units to Fennoscandia. In Fig. 3 a, 210 the early-Sveconorwegian closure of an ocean between the Telemarkia and Idefjorden 211 lithotectonic units resulted in the formation of the Bamble-Kongsberg lithotectonic units at c. 212 1150–1120 Ma, and accretion of an exotic Telemarkia lithotectonic unit (Bingen et al., 2005). 213 In Fig. 3 b, the five lithotectonic units are all endemic to Fennoscandia. Only large-scale 214 movements between them are considered, steered by large scale wrench tectonics, which is 215 controlled by strike-slip shearing along the main Sveconorwegian shear zones (Bingen et al., 216 2005; Lamminen and Köykkä, 2010; Stephens and Wahlgren, 2020b). In Fig. 3 c, closure of 217 an oceanic basin at c. 990 Ma between the Eastern Segment and the Idefjorden lithotectonic
218 unit along the Mylonite Zone, resulted in collision between Baltica (Fennoscandia) and a
219 continent composed of the four western lithotectonic units of the orogen ('Sveconorwegia')
220 (Möller and Andersson, 2018; Petersson et al., 2015b). In Fig. 3 d, an (Himalaya-Tibet type)
221 collision at and after c. 1065 Ma between Baltica (Fennoscandia) and one (or several)
222 continental plate(s) (possibly Amazonia, Laurentia, and intervening terranes exposed in
223 Mexico and the Andes of Colombia) involved closure of oceanic basins to the west of the
224 exposed orogen (Bingen et al., 2008c; Bogdanova et al., 2008; Cawood and Pisarevsky, 2017;
225 Ibanez-Mejia et al., 2011; Stephens and Wahlgren, 2020b; Weber et al., 2010). In Fig. 3 e and
226 f, non-collisional (Andean type) models feature an eastward subduction of an oceanic plate
227 below the western margin of Baltica (Fennoscandia) during the entire Sveconorwegian
228 orogeny, from 1150 to 900 Ma, in the absence of a final collision. The lithotectonic units in
229 the orogen either were assembled already before the Sveconorwegian orogeny (Fig. 3 e)
230 (Falkum and Petersen, 1980; Slagstad et al., 2013) or, alternatively, they were dispersed
231 during the pre-Sveconorwegian time interval (1280–1150 Ma) and then re-assembled during
232 the Sveconorwegian orogeny after 1150 Ma (Fig. 3 f) (Slagstad et al., 2020). These six
233 models are not mutually exclusive because terrane assembly (Fig. 3 a, c) can be anticipated
234 before a collision (Fig. 3 d) or during a protracted subduction history (Fig. 3 e, f), and because
235 deformation partitioning (Fig. 3 b) can take place before, during and after a collision or during
236 protracted subduction. Arguments supporting or dismissing aspects of each of these orogenic
237 models are discussed in more detail below.

238 **2.4 Secular evolution of the Earth, mantle delamination and orogenic plateau**

239 Estimates of heat flow and heat production through Earth history suggest that the
240 asthenosphere was c. 100°C hotter in the Mesoproterozoic than at present (Gerya, 2014;
241 Herzberg et al., 2010; Johnson et al., 2013; Korenaga, 2008; Sizova et al., 2014). A hotter
242 asthenosphere implies a weaker rheology of lithospheric plates. The tectonic consequences of
a hotter asthenosphere on the dynamics of orogeny are multiple and include, but are not
limited to, ductile thick-skinned deformation, lower topography, more proximal
sedimentation, shallower slab breakoff, widespread partial melting in the lower to middle
crust, widespread syn-orogenic magmatism, ultrahigh temperature granulite-facies
metamorphism (above 900°C), decoupling between crust and lithospheric mantle, and
remelting of basaltic underplates to produce anorthosite plutons (Brown, 2006, 2013; Gerya,
2014; Rey and Houseman, 2006; Sizova et al., 2014; Vander Auwera et al., 2011;
Vanderhaeghe, 2012). These consequences can be evaluated qualitatively in the Proterozoic
geological record (Cagnard et al., 2011; Chardon et al., 2009). However, they are difficult to
assess and quantify individually (Sizova et al., 2014).

There is wide consensus that after the Archean, plate tectonics has imposed dominant
horizontal movements to orogenies. However, an evaluation of the composition and
temperature of the lithosphere through Earth history suggests that, after the Archean, the sub-
continental lithospheric mantle was, on average, denser than the asthenosphere (Griffin et al.,
2009; Poudjom Djomani et al., 2001). The sub-continental lithospheric mantle was therefore
gravitationally unstable in the Proterozoic, like in the Phanerozoic, and prone to delamination
and foundering (subduction) (Bird, 1979; Chen et al., 2017; Krystopowicz and Currie, 2013).
Delamination of the lithospheric mantle is compensated by upwelling of asthenosphere. The
parameters and geometry of delamination in convergent orogens were explored numerically
by Li et al. (2016). Delamination is promoted by the density contrast between the lithospheric
mantle and the asthenosphere, rheological weakness of the lower crust and the lithospheric
mantle, convergence rate, and eclogitization of the lower crust.

The Sveconorwegian orogeny is characterized by widespread crustal partial melting, low-
pressure-ultrahigh-temperature metamorphism and massif-type anorthosite plutonism, typical
of hot orogens. Orogenic models for the Sveconorwegian orogeny should integrate the
268 evolution of the mantle (and not only the crust). They require consideration of vertical
269 movements of the lithospheric mantle and the asthenosphere in addition to horizontal
270 movements of the lithospheric plates.

271 Orogenic plateaus are a hallmark of large and hot convergent orogens (Beaumont et al.,
272 2006; Godin et al., 2006; Jamieson and Beaumont, 2013; Li et al., 2016; Rey et al., 2001;
273 Royden et al., 2008; Vanderhaeghe, 2012). An orogenic plateau consists of elevated and
274 thickened crust spreading by gravitational forces, above a lithospheric mantle thinned by
275 delamination. Temperature in the crust is regulated by self-heating and basal heating from the
276 mantle. The crust of a plateau is characterized by a little viscous low- to middle-crust,
277 weakened by partial melting, called infrastructure, overlain by a brittle upper crust, called
278 superstructure or orogenic lid (Jamieson and Beaumont, 2013; Rey et al., 2001;
279 Vanderhaeghe, 2012). In the infrastructure, metamorphism typically carries a high-
280 temperature signature, overprinting pre-plateau metamorphic signatures (for example early
281 high-pressure metamorphism) (Godin et al., 2006). Due to the difference in viscosity, the
282 infrastructure and superstructure are structurally decoupled. The infrastructure can flow under
283 the superstructure (channel flow), leading to a situation where the superstructure is in
284 extension, while the infrastructure is in compression. An orogenic plateau can be anticipated
to grow with time if convergence is maintained (Li et al., 2016; Royden et al., 2008).
285
286 The Sveconorwegian orogen consists of a patchwork of high-grade gneiss complexes and
287 low-grade rocks (Fig. 4). In this paper, these are interpreted as the remnants of the
288 infrastructure and superstructure of an orogenic plateau, respectively.

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**Figure 4.** Sketch map of the Sveconorwegian orogen, showing the extent of the infrastructure and superstructure (orogenic lid) of the orogen during the main Sveconorwegian orogeny (1065–920 Ma) and Sveconorwegian plutons (1065–920 Ma).

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New samples: Eidsvoll: B0935, B0936, B0940, B0941; Auli: B0937, B0938, B0939

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**Figure 5.** Sketch map of the Sveconorwegian orogen, with emphasis on Sveconorwegian events younger than c. 1100 Ma. Localities of samples in Auli and Eidsvoll analysed in this study are shown. Age intervals in the legend rounded in 5 Ma intervals.

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<figure><img src="image_9.png" /><figcaption>**Figure 6.** Sketch map of the Sveconorwegian orogen, with emphasis on pre- and early-Sveconorwegian events and rocks.</figcaption></figure>

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**Figure 7.** Kernel density estimators summarizing the geochronology of magmatic rocks in the five lithotectonic units of the Sveconorwegian orogen, on a compilation of published data.

The plots are generated with "DensityPlotter" (Vermeesch, 2012) (each published age is entered as one value, with a bandwidth of 6 Ma; the height of the five curves is identical and normalized to the one of the largest peak). The compilation is provided in the supplementary material, with referencing.

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<figure><img src="image_11.png" /></figure>

**Figure 8.** Hafnium isotopic composition of magmatic rocks in the Sveconorwegian orogen and Fennoscandia foreland, expressed as εHf (initial value) as a function of intrusion age. Interpretations of the distribution of data are discussed in the text. The five lithotectonic units are shown with distinct colours and summarized by an evolution vector. The Oaxaquia lithotectonic unit (Mexico) and the inliers in the Andes of Colombia are shown for comparison. Each symbol represents the average value for one sample, of the isotopic composition of several analyses of zircon or baddeleyite or of one whole-rock analysis (only a few samples), at the recommended time of intrusion (zircon or baddeleyite U–Pb age). Total of 220 samples. Sources of data: Sveconorwegian orogen and foreland: Andersen et al. (2009; 2002b; 2007), Lamminen et al. (2011), Pedersen et al. (2009), Petersson et al. (2015a; 2015b), Roberts et al. (2013); Söderlund et al. (2005); Oaxaquia and inliers in the Andes of Colombia: Ibanez-Mejia et al. (2015), Weber et al. (2010); DM: depleted mantle (Griffin et al., 2000); OVAM: oceanic volcanic arc mantle (Dhuime et al., 2011); CHUR: chondritic reservoir (Bouvier et al., 2008). The top right inset shows the ¹⁷⁶Lu/¹⁷⁷Hf ratio and evolution vectors of isotopic reservoirs and typical zircon.

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### 3 Geology of the Sveconorwegian orogen
The geology of the Sveconorwegian orogen is reviewed below from east to west, using the nomenclature summarized in Table 1 and the maps of Fig. 2, Fig. 4, Fig. 5 and Fig. 6. A compilation of the geochronology of magmatic rocks is provided in Fig. 7, and a compilation of Lu-Hf isotopic data in Fig. 8.

## 3.1 Fennoscandian foreland

The Fennoscandian foreland of the Sveconorwegian orogen (Fig. 2) comprises mainly metamorphosed Paleoproterozoic magmatic rocks (plutonic and volcanic rocks) and siliciclastic sedimentary rocks, dating back to between c. 1960 Ma and 1740 Ma (Bergman et al., 2008; Korja et al., 2006; Lahtinen et al., 2009; Stephens, 2020). These rocks were assembled during the accretionary Svecokarelian orogeny. They were unconformably overlain and crosscut by post-Svecokarelian volcanic and plutonic complexes formed between c. 1710 and 1680 Ma (Appelquist et al., 2011; Brander et al., 2012; Högdahl et al., 2004; Ripa and Stephens, 2020a). These rocks are attributed in the literature to the Phase 2 of the Transcandinavian Igneous Belt and are little deformed to undeformed. They are interpreted to have formed in a supra-subduction geodynamic setting after the Svecokarelian orogeny.

Younger Mesoproterozoic magmatic rocks intruded this basement, including granite plutons (1530–1220 Ma; Andersson et al., 2002b; Brander and Söderlund, 2009; Cecys and Benn, 2007; Johansson et al., 2016), dolerites (c. 1460 Ma; Söderlund et al., 2005), and the so-called Central Scandinavian dolerites (1271 ± 1 to 1246 ± 2 Ma; Brander et al., 2011; Ripa and Stephens, 2020c; Söderlund et al., 2006).

The c. 1 km thick unconformable Jotnian sandstone was deposited in a gentle continental sag basin between c. 1580 Ma and 1270 Ma and it is not deformed (Lundmark and Lamminen, 2016; Ripa and Stephens, 2020b). The southernmost part of the Fennoscandian foreland was reworked during the Hallandian orogenic event between 1465 and 1385 Ma
354 (Fig. 6) (Bogdanova et al., 2008; Brander and Söderlund, 2009; Ulmius et al., 2015; Wahlgren
355 and Stephens, 2020).

356 Sveconorwegian-related brittle deformation reached far into the Fennoscandian foreland
357 (Andréasson and Rodhe, 1994; Elminen et al., 2018; Mattila and Viola, 2014; Saintot et al.,
358 2011; Viola et al., 2009; Viola et al., 2013). The Blekinge-Dalarna dolerites form a weakly
359 arcuate N-S trending dyke swarm parallel to the Sveconorwegian front (Fig. 4; Fig. 5). They
360 intruded between 978 ± 2 and 946 ± 1 Ma, in the easternmost part of the Sveconorwegian
361 orogen and its foreland (Gong et al., 2018; Ripa and Stephens, 2020d; Söderlund et al., 2005).
362 The c. 1200 m thick, sandstone dominated, Almesåkra Group represents possible remnants of
363 a Sveconorwegian fold-and-thrust belt, to the east of the Sveconorwegian front (Fig. 5) (Ripa
364 and Stephens, 2020d; Rodhe, 1987). Locally preserved peperitic contacts between the
365 Blekinge-Dalarna dolerites and these sediments suggest that the sandstone was
366 unconsolidated during intrusion of the dolerites and therefore that the two rock types are
367 broadly coeval.

368 The Neoproterozoic, c. 1400 m thick, microfossil-bearing, Visingsö Group is exposed
369 along the Sveconorwegian front in Sweden (Fig. 5). Its deposition is bracketed between 886 ±
370 9 Ma (detrital zircon U–Pb data) and c. 740 Ma (biostratigraphy). It can be considered as the
371 infill of a post-Sveconorwegian, fault-controlled basin (Loron and Moczydłowska, 2018;
372 Moczydłowska et al., 2018; Pulsipher and Dehler, 2019; Wickström and Stephens, 2020).

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**Figure 9.** Microphotographs of thin sections showing the contrast between high-pressure (eclogite-facies) and ultrahigh temperature (granulite-facies) metamorphism, east and west of the Sveconorwegian orogen, respectively, at c. 1000 Ma. (a) Kyanite-bearing (retro)eclogite from the eclogite-bearing nappe in the Eastern Segment (Möller and Andersson, 2018). The thin section shows a partly preserved peak eclogite-facies assemblage of garnet (Grt) + omphacite (Cpx) + kyanite (Ky) + amphibole + rutile (1.8 GPa - 870°C - 988 ± 6 Ma) breaking down into a symplectitic (sy) assemblage during isothermal decompression.

Symplectites (sy) include a sapphirine + corundum + anorthite reaction rim around kyanite, an orthopyroxene + plagioclase + amphibole reaction rim around clinopyroxene, and plagioclase expulsion symplectite in former omphacite (Cpx). Garnet preserves a prograde (pre-eclogite-facies) zoning, and the rock shows evidence for a hairpin P-T path (Tual et al., 2017). (b)

Sapphirine + orthopyroxene granulite from the Ivesdal locality, in the ultra-high temperature (UHT) zone of Rogaland, in the Telemarkia lithotectonic unit (Laurent et al., 2018b). The thin section shows the peak assemblage of sapphirine (Spr) mantled by orthopyroxene (Opx) (0.6 GPa - 920°C - 1029 ± 9 to 1006 ± 8 Ma) breaking down into an assemblage of cordierite (Crd) + hercynite (Sp) with additional biotite (Bt) (4.5 GPa – 900°C) giving evidence for a clockwise P-T path.

-----------------

**3.2 Eastern Segment**
394 **3.2.1 Svecokarelian and post-Svecokarelian evolution**

395 The Eastern Segment is a 60 to 120 km wide, N–S trending lithotectonic unit mainly
396 consisting of granitic to quartz-monzonitic orthogneiss (Fig. 2) (Berthelsen, 1980; Möller and
397 Andersson, 2018; Stephens and Wahlgren, 2020a). The protoliths formed between c. 1900
398 and 1660 Ma, with a strong frequency maximum of crystallization ages between 1710 and
399 1660 Ma (Fig. 7). They have an alkali-calcic geochemical composition and are characterized
400 by a mildly positive εHf and εNd isotopic signature (average εHf = + 3.0 at 1700 Ma; Fig. 8)
401 (Appelquist et al., 2011; Appelquist et al., 2008; Brander et al., 2012; Gorbatschev and
402 Bogdanova, 2006; Petersson et al., 2015a; Söderlund et al., 1999; Söderlund et al., 2002;
403 Stephens and Wahlgren, 2020a). They represent the western continuation of the
404 Paleoproterozoic crust exposed in the foreland of the Sveconorwegian orogen, especially the
405 post-Svecokarelian, 1710–1680 Ma, magmatic rocks exposed just east of the Sveconorwegian
406 front (Petersson et al., 2015a; Ripa and Stephens, 2020a; Stephens and Wahlgren, 2020a).
407 They were presumably formed in the same geodynamic setting along the same active
408 continental margin.

409 **3.2.2 Hallandian and pre-Sveconorwegian evolution**

410 After an event of mafic magmatism around 1565 Ma (Beckman et al., 2017; Söderlund et al.,
411 2004; Söderlund et al., 2005), the southern part of the Eastern Segment and the
412 Sveconorwegian foreland were together affected by the Hallandian orogeny (Fig. 6; Fig. 7).
413 The Hallandian orogeny involved low-pressure amphibolite- to granulite-facies
414 metamorphism, migmatitization and deformation between c. 1465 and 1385 Ma (Brander et
415 al., 2012; Möller et al., 2007; Piñán-Llamas et al., 2015; Söderlund et al., 2002; Ulmius et al.,
416 2015), and was accompanied by magmatism during the same time interval (Fig. 7) (Åhäll et
417 al., 1997; Andersson et al., 1999; Brander and Söderlund, 2009; Cecys et al., 2002;
418 Christoffel et al., 1999; Möller et al., 2015; Ulmius et al., 2015). The final stage of Hallandian
419 magmatism includes a suite of charnockite-mangerite, granite and anorthosite plutons formed
420 between c. 1400 and 1380 Ma (Åhäll et al., 1997; Christoffel et al., 1999; Harlov et al., 2013;
421 Möller et al., 2015). The Hallandian orogeny may have involved subduction along the
422 southern margin of Baltica and may record a change in the configuration of subduction zones
423 around Baltica, from E-dipping before 1480 Ma to N-dipping after 1465 Ma (Pisarevsky et
424 al., 2014; Roberts and Slagstad, 2015; Stephens and Wahlgren, 2020b; Ulmius et al., 2015).

425 Post-Hallandian bimodal plutonism took place between 1225 and 1180 Ma, including
426 dolerites (Protogine zone dolerites) and syenitic to granitic plutons (e.g. the Vaggeryd syenite;
427 Fig. 6) (Larsson and Söderlund, 2005; Petersson et al., 2015a; Söderlund and Ask, 2006;
428 Söderlund et al., 2005). These rocks are characterized by a supra-chondritic (radiogenic) εHf
429 isotopic signature (+1.2 < εHf < + 6.6) implying an influx of depleted mantle derived magmas
430 along the Sveconorwegian front (Fig. 8) (Petersson et al., 2015a; Söderlund et al., 2005).

431 **3.2.3 Sveconorwegian orogeny**

432 The Sveconorwegian metamorphic grade in the Eastern Segment increases towards the WSW
433 (Fig. 5) (Johansson et al., 1991; Möller and Andersson, 2018; Möller et al., 2015; Piñán-
434 Llamas et al., 2015). Four zones of distinct metamorphic and structural reworking can be
435 defined from east to west: (i) a frontal wedge, (ii) a transitional section, (iii) an internal
436 section and (iv) an eclogite-bearing ductile nappe (Möller and Andersson, 2018; Möller et al.,
437 2015).

438 The frontal wedge (i) is a zone of non-penetrative Sveconorwegian deformation forming a
439 steep or fan-shaped structure in cross section that narrows and steepens towards the south
440 (Möller and Andersson, 2018; Stephens and Wahlgren, 2020a; Wahlgren et al., 1994). The
441 zone comprises a network of thin (<100 m), N–S trending, steeply dipping, greenschist- to
442 amphibolite-facies ductile shear zones with mainly western-block-up kinematics (Andréasson
443 and Dallmeyer, 1995; Brander et al., 2012; Gorbatschev and Bogdanova, 2006; Söderlund et
al., 2004; Wahlgren et al., 1994). The frontal wedge is bound in the east by the

Sveconorwegian front, which in the north is a system of discontinuous west dipping shear

zones with reverse top-to-east sense of shear (Wahlgren et al., 1994). In the northernmost part

of the Eastern Segment in Norway, the frontal wedge is poorly documented.

The transitional section (ii) exhibits a near-penetrative amphibolite-facies overprint, with
little evidence for partial melting (Beckman et al., 2017; Möller and Andersson, 2018;

Söderlund et al., 1999). The internal section (iii) is characterized by upper-amphibolite-facies
conditions increasing westwards to high-pressure granulite-facies conditions (1.1 GPa -

850°C; Fig. 5). This metamorphic evolution caused widespread migmatitization, transposition

leading to mafic and felsic gneissic layering (banding), dynamic recrystallization of original
magmatic textures, as well as reworking of previous Hallandian structures, where present

(Andersson et al., 1999; Connelly et al., 1996; Hansen et al., 2015; Möller et al., 2015; Möller
et al., 2007; Piñán-Llamas et al., 2015). The regional aeromagnetic map (Geological Survey

of Sweden) unveils prominent, regional scale fold interference patterns, with E-W trending
and gently-plunging fold axes and trains of N-S trending folds (Möller et al., 2007; Stephens
and Wahlgren, 2020a; Viola et al., 2011). Several generations of folds can be recognised (F1
to F4), including km-scale asymmetric to recumbent folds and late upright folds (Möller and

Andersson, 2018; Möller et al., 2015; Piñán-Llamas et al., 2015; Tual et al., 2015). These

different generations record continued deformation under high-grade metamorphic conditions.

The eclogite-bearing ductile nappe (iv) is hosted in the innermost section of the Eastern

Segment as an E-vergent ductile nappe, folded into a c. 50 x 75 km large recumbent fold (Fig.

5) (Möller et al., 2015; Tual et al., 2015). It is well defined on the regional aeromagnetic map
and bordered (on the southern and eastern flanks) by a sheet of c. 1380 Ma granite (Fig. 6).

The ductile nappe hosts retro-eclogite bodies up to 2 km in length (Möller, 1998, 1999;

Möller and Andersson, 2018; Möller et al., 2015; Tual et al., 2015). The retro-eclogite bodies
469 are layered mafic rocks, including two characteristic varieties, a Mg-Al-rich kyanite-bearing
470 variety and a Fe-Ti-rich variety. Retro-eclogites preserve prograde growth zoning of garnet
471 and show widespread retrogression of omphacite and kyanite into granulite-facies
472 symplectites (clinopyroxene + plagioclase, orthopyroxene + plagioclase, and anorthite +
473 sapphirine + corundum) (Fig. 9 a). They constrain a narrow (hairpin) clockwise pressure-
474 temperature path at high temperature. Eclogite-facies peak conditions of 1.65–1.9 GPa and
475 850–900°C were followed by near-isothermal decompression (Tual et al., 2017). Eclogite
476 boudins are hosted in strongly deformed, partly migmatitic, gneisses characterized by a
477 pervasive foliation, E–W stretching lineation, and S- to E-vergent folds (Möller et al., 2015;
478 Tual et al., 2015). Two post-eclogite-facies deformation phases (D1-D2) are described as
479 successively documenting an early stage of exhumation and east-directed transport of the
480 ductile nappe, lubricated by partial melts.

481 A zircon U–Pb age determination from an eclogite sample defines a maximum age for
482 eclogite-facies metamorphism of 988 ± 6 Ma (Möller et al., 2015). Zircon in felsic and mafic
483 gneiss, migmatite and syn-kinematic granite in the entire Eastern Segment, including the
484 eclogite-bearing nappe, yields a consistent age interval between 978 ± 7 and 961 ± 6 Ma for
485 amphibolite- to granulite-facies metamorphism, deformation and partial melting (Andersson
486 et al., 2002a; Beckman et al., 2017; Hansen et al., 2015; Möller et al., 2015; Möller et al.,
487 2007; Piñán-Llamas et al., 2015; Söderlund et al., 2002). Cross-cutting pegmatite dykes
488 intruded between 961 ± 13 and 934 ± 6 Ma (Andersson et al., 1999; Möller et al., 2007;
489 Möller and Söderlund, 1997; Söderlund et al., 2008b; Söderlund et al., 2002). Titanite U–Pb
490 ages range from c. 976 ± 4 to 923 ± 3 Ma, with the oldest age recorded in the northern part of
491 the transitional section and the youngest ages in the internal section (Connelly et al., 1996;
492 Johansson et al., 2001; Söderlund et al., 1999; Wang et al., 1998). Hornblende and biotite
493 $^{40}$Ar/$^{39}$Ar plateau ages in the internal section are interpreted to date regional cooling between
c. 530 and 330°C between c. 901 ± 2 and 893 ± 3 Ma (Ulmius et al., 2018). Biotite and muscovite <sup>40</sup>Ar/<sup>39</sup>Ar plateau ages collected in the frontal wedge range from 930 ± 6 to 882 ± 2 Ma (Andréasson and Dallmeyer, 1995; Page et al., 1996a; Ulmius et al., 2018). The youngest ages are recorded in the southernmost exposed section of the orogen. These ages either record discrete events of (re)crystallization or cooling after deformation by shear zones at the front of the orogen (Andréasson and Dallmeyer, 1995; Page et al., 1996a; Ulmius et al., 2018).

500

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**Figure 10.** New geochronological data of migmatitic gneisses in the Eidsvoll-Auli area, Idefjorden lithotectonic unit (Fig. 5). (a–k) Tera-Wasserburg concordia diagrams with zircon SIMS U–Pb analyses and a selection of CL images of zircon with position of analyses. Blue ellipses for magmatic zircon cores, green ellipses for detrital zircon cores, and red ellipses for low-U sector zoned zircon and zircon rims attributed to migmatitization. One sigma error ellipses. (e) Photo of outcrop where two samples represent two generations of leucosomes, with B0938 crosscutting B0937. Interpretation: migmatites from the five studied localities are characterized by abundant, interconnected leucosomes (stromatic texture) parallel to the gneissic foliation. Zircon contains an inherited core (magmatic or detrital), a CL-dark mantle and a CL-bright rim. Analyses of the mantle overlap with those of the core and define a significant spread in each sample. The spread of apparent ages can be interpreted to represent partial recrystallization of the core during partial melting. Newly formed CL-bright rims or large crystals with oscillatory to weakly sector zoning reflect crystallization of zircon related to migmatitization (Harley et al., 2007; Kelsey et al., 2008; Rubatto et al., 2009) between 1039 ± 17 and 997 ± 16 Ma (a, c, g, i, k). In this c. 40 Myr time interval, the biotite-muscovite and biotite-garnet leucosomes range from 1039 ± 17 to 1022 ± 23 Ma (d, g, i, k), while the hornblende-biotite-bearing leucosomes are marginally to significantly younger with ages of 1003 ± 34 and 997 ± 16 Ma (a, c). This difference suggests that muscovite and biotite dehydration melting took place before amphibole dehydration melting, in what can be interpreted as reflecting increasing temperature or isothermal decompression.

**Table 2.** Summary of sampling and zircon U-Pb data for migmatitic gneisses, Eidsvoll-Auli area, Idefjorden lithotectonic unit.

**3.3 Idefjorden lithotectonic unit**
527 The Idefjorden lithotectonic unit is a c. 140 km wide unit exposed west of the Eastern Segment on either side of the Permian Oslo Rift (Fig. 2, Fig. 5) (Åhäll and Connelly, 2008; Åhäll and Gower, 1997; Bergström et al., 2020; Bingen et al., 2001; Park et al., 1991; Viola et al., 2011). It is bounded in the east by the 450 km long, west dipping, Mylonite Zone.

### 3.3.1 Gothian and pre-Sveconorwegian evolution

532 The Idefjorden lithotectonic unit is made up of plutonic and volcanic rocks formed during the Gothian accretionary orogeny mainly between 1660 and 1520 Ma and associated with metasedimentary rocks (Fig. 7) (Åhäll and Connelly, 2008; Åhäll and Larson, 2000; Ahlin et al., 2006; Andersen et al., 2004a; Bergström et al., 2020; Bingen et al., 2005; Brewer et al., 1998; Graversen and Pedersen, 1999). From east to west, three complexes (called formations or belts in the literature) are described as younging towards the west (Åhäll and Connelly, 2008; Brewer et al., 1998): (i) the 1660–1640 Ma metavolcanic Horred Complex, (ii) the 1630–1590 Ma metavolcanic and metasedimentary Åmål Complex associated with the Göteborg granite suite, and (iii) the 1590–1520 Ma metasedimentary and metavolcanic Stora Le-Marstrand Complex, associated with the 1580–1520 Ma plutonic Hisingen Suite. The Stora Le-Marstrand Complex, exposed east of the Oslo Rift, correlates with the Veme Complex west of the Oslo Rift (Fig. 6) (Bingen et al., 2001). The Stora Le-Marstrand and Veme complexes comprise several metasedimentary successions (Åhäll and Connelly, 2008), consisting of thick packages of turbiditic psammite and greywacke metamorphosed under amphibolite-facies conditions (Bingen et al., 2001). Sedimentation started before c. 1585 Ma (metagreywacke xenoliths in a $1584 \pm 7$ Ma granite pluton) and continued to after c. 1500 Ma (detrital zircon geochronology in 12 samples) (Åhäll and Connelly, 2008; Åhäll et al., 1998; Andersen et al., 2004a; Bingen et al., 2001; Bingen and Viola, 2018). The paragneisses analysed in this study just east of the Oslo Rift (Eidsvoll and Auli; Fig. 5; Fig. 6; Fig. 10) are attributed to the Stora Le-Marstrand Complex.
The c. 1660–1520 Ma (Gothian) magmatic suites (Fig. 7) are characterized by low- to medium-K calc-alkaline geochemical compositions, with supra-chondritic Hf and Nd isotopic signature (average εHf = +4.8 in the Idefjorden lithotectonic unit at 1570 Ma; Fig. 8), reflecting continental and oceanic volcanic arc magmatism (Andersen et al., 2004a; Andersen et al., 2002b; Bergström et al., 2020; Brewer et al., 1998; Petersson et al., 2015b). Metabasalts interlayered in the Stora Le-Marstrand rocks are tholeiitic and interpreted as oceanic back-arc magmatism (Brewer et al., 1998).

The c. 1660–1520 Ma rocks were assembled during the Gothian accretionary orogenic event (Åhäll and Connelly, 2008; Andersen et al., 2004a; Petersson et al., 2015b; Roberts and Slagstad, 2015). Convincing evidence for Gothian regional deformation and metamorphism includes crosscutting relationships (folded xenoliths in a 1584 ± 7 Ma pluton) and U–Pb geochronological data in zircon and monazite ranging from 1546 ± 5 to 1539 ± 8 Ma from a few localities in the Veme and Stora Le-Marstrand complexes (Åhäll and Connelly, 1998, 2008; Bingen et al., 2008b; Bingen and Viola, 2018; Connelly and Åhäll, 1996).

The 1660–1520 Ma rocks are intruded by the 1457 ± 6 Ma, N–S trending Orust tholeiitic dolerite dykes (Åhäll and Connelly, 1998), and the 1340–1305 Ma bimodal Kungsbacka suite (Fig. 6) (Austin Hegardt et al., 2007). The Dal Group (or Dalsland Group) is a c. 2 km thick succession of low-grade clastic sedimentary rocks and tholeiitic basalt, exposed in a syncline structure, overlying (and therefore younger than) the Kungsbacka suite (Fig. 6) (Brewer et al., 2002). The Dal Group is poorly characterized. However, it may provide critical evidence for the tectonic evolution of the Idefjorden lithotectonic unit before the Sveconorwegian orogeny (Brewer et al., 2002) and therefore would warrant new investigations.

### 3.3.2 Sveconorwegian orogeny

In the Idefjorden lithotectonic unit, the Sveconorwegian deformation is associated with a N–S to NW–SE structural grain and has variable strain intensity. Several shear zones, including the
prominent Ørje and Göta Älv shear zones (Fig. 2), are parallel to this structural grain (Bergström et al., 2020; Park et al., 1991; Viola et al., 2011; Wahlgren et al., 2015).

Metamorphism ranges from greenschist- to granulite-facies. The low-grade rocks are exposed in syncline structures (Fig. 4). For example, between the Göta Älv shear zone and the Mylonite Zone, the Gillberga syncline hosts the Glaskogen low-grade complex, bounded by low-angle shear zones (Lindh et al., 1998), and the Åmal volcanic rocks (1614 ± 7 Ma), known for good preservation of primary volcanic structures (Lundqvist and Skiöld, 1993). In the amphibolite-facies gneiss complexes east and west of the Göta-Älv shear zone, garnet amphibolites provide pressure-temperatures estimates of 0.9 to 1.2 GPa – 730 to 790 °C (3 samples; Austin Hegardt, 2010). High-pressure garnet–clinopyroxene-bearing granulite-facies assemblages are reported from metadolerite dykes hosted in amphibolite-facies gneisses from several localities east of the Göta-Älv shear zone (Trollhättan, Fig. 5) (Söderlund et al., 2008a). Geothermobarometry coupled with zircon U–Pb data and mineral isochron data from two dykes indicate conditions of c. 1.5 GPa – 740 °C at 1046 ± 6 Ma and c. 1.0 GPa – 700 °C at 1026 ± 5 Ma (Söderlund et al., 2008a). In the Veme Complex, west of the Oslo rift in Norway, a kyanite-garnet-rutile paragneiss hosting clinopyroxene-garnet-plagioclase-rutile mafic boudins also records high-pressure granulite-facies conditions, with pressure-temperature estimates of 1.2 GPa – 780 °C (Hensmoen, Fig. 5) (Bingen et al., 2008b). Monazite in the kyanite-rutile-gneiss records peak metamorphism at 1052 ± 4 Ma (Bingen et al., 2008b).

East of the Oslo rift in Norway, amphibolite-facies metamorphism is associated with a foliation dipping unimodally to the NE and with folds verging to the W to SW (Viola et al., 2011). The timing of this metamorphism is provided by the new U–Pb data from zircon rims in migmatitic samples (Eidsvoll–Auli area, Fig. 5, Fig. 10, Table 2). The dates range from 1039 ± 17 to 997 ± 16 Ma, in seven samples affected by both muscovite- biotite- and
602 amphibole-dehydration melting. The dates are interpreted to record crystallization of the
603 leucosomes. This interval overlaps with published zircon and titanite U-Pb data and a Sm-Nd
604 mineral isochron interpreted to record high-grade metamorphism between 1043 ± 11 Ma and
605 1024 ± 9 (7 samples ; Åhäll et al., 1998; Austin Hegardt, 2010; Austin Hegardt et al., 2007;
606 Bingen et al., 2008b), and also with intrusion of rare-mineral pegmatites between 1041 ± 2
607 and 984 ± 6 Ma (Romer and Smeds, 1996).

608 Several mafic to felsic magmatic intrusions, with a consistent WNW-ESE trend and dated
609 between 951 ± 7 and 915 ± 1 Ma, crosscut the regional amphibolite-facies ductile fabric in the
610 coastal area of Sweden (Årebäck et al., 2008; Hellström et al., 2004; Scherstén et al., 2000;
611 Wahlgren et al., 2015). These include a lamprophyre dyke (915 ± 1 Ma) (Wahlgren et al.,
612 2015) and the small Hakefjorden norite-anorthosite complex (916 ± 11 Ma), carrying
613 evidence for extensive fractional crystallization (Årebäck and Stigh, 2000). The Flå and
614 Bohus biotite-muscovite granite plutons intruded between 932 ± 8 and 922 ± 3 Ma, as large
615 tabular bodies, in pressure conditions of c. 0.4 GPa (Fig. 5) (Eliasson et al., 2003; Eliasson
616 and Schöberg, 1991; Lamminen et al., 2011). A final batch of rare-mineral pegmatite formed
617 between 909 ± 1 and 906 ± 6 Ma (Müller et al., 2017).

3.3.3 *The Mylonite Zone*

619 The Mylonite Zone is a generally west dipping shear zone juxtaposing the Eastern Segment
620 and Idefjorden lithotectonic unit. It is several km thick, continuous for some 450 km and
621 characterized by a widespread greenschist- to upper amphibolite-facies mylonitic fabric (Fig.
622 2, Fig. 5) (Andersson et al., 2002a; Bergström et al., 2020; Möller et al., 2015; Park et al.,
623 1991; Stephens et al., 1996; Viola and Henderson, 2010; Viola et al., 2011). It possibly roots
624 in the mantle (EUGENO-S-working-group, 1988).

625 The Mylonite Zone is interpreted as a Sveconorwegian mid-crustal thrust zone placing the
626 Idefjorden lithotectonic unit on top of the Eastern Segment, with an overall southeastward
transport direction oblique to the orogen (Stephens et al., 1996; Viola and Henderson, 2010; Viola et al., 2011). Shear zones inside the Idefjorden lithotectonic unit, including the Ørje and Göta Älv shear zones (Fig. 2) are similarly interpreted as transpressional thrust zones (Park et al., 1991; Viola et al., 2011; Wahlgren et al., 2015).

In the north (in Norway), the Mylonite Zone trends NW-SE and has a steep attitude with sinistral strike-slip kinematics. This segment has been interpreted as the sinistral lateral ramp to the thrust frontal ramp farther to the southeast. The frontal ramp dips gently to moderately to the west and bears a NW plunging stretching lineation associated with dominant top-to-southeast reverse displacement. In the southernmost part, the shear zone turns quite abruptly E-W, dipping gently to the north, and accommodating a dominant component of dextral strike-slip shearing. This part is interpreted as a dextral lateral ramp of the thrust zone (Viola and Henderson, 2010; Viola et al., 2011). The importance of the southernmost dextral lateral ramp is downplayed by Bergström et al. (2020), who interpret the Mylonite Zone, as a whole, as a sinistral transpressional thrust zone. Zircon U-Pb data in the Mylonite Zone and close hanging wall and footwall record amphibolite-facies migmatitization and associated ductile deformation between 980 ± 13 and 969 ± 13 Ma (Andersson et al., 2002a).

The Mylonite Zone was reactivated in extension with top-to-the-west kinematics along a network of localized shear zones, contributing to exhumation of the Eastern Segment in the footwall (Viola and Henderson, 2010; Viola et al., 2011). Muscovite and biotite <sup>40</sup>Ar/<sup>39</sup>Ar data suggest that this deformation took place between 923 ± 4 and 861 ± 5 Ma (Viola et al., 2011).

## 3.4 Kongsberg and Bamble lithotectonic units

The Bamble and Kongsberg lithotectonic units are two narrow c. 25 km wide units situated in the center of the exposed Sveconorwegian orogen (Fig. 2). Kongsberg trends N-S while Bamble trends NE-SW. These two lithotectonic units share a number of features, including
652 evidence for early-Sveconorwegian metamorphism (1150-1120 Ma) (Bingen et al., 2008b;
653 Bingen and Viola, 2018; Engvik et al., 2016; Knudsen et al., 1997; Nijland et al., 2014;
654 Starmer, 1985; Viola et al., 2016).

655 3.4.1 Gothian–Telemarkian evolution

656 Two main lithological complexes are present in the Bamble and Kongsberg lithotectonic
657 units, (i) an orthogneiss complex, referred to as Kongsberg Complex in Kongsberg and
658 Bamble Complex in Bamble, and (ii) a quartzite-dominated metasedimentary complex, called
659 Modum Complex in Kongsberg and Nidelva and Kragerø Complexes in Bamble (Fig. 6). (i)
660 The orthogneiss complex consists of penetratively deformed orthogneisses with composition
661 ranging from dioritic to tonalitic, to granitic, and more competent gabbro plutons (Holleia and
662 Blengsvatn; Bingen and Viola, 2018; Nijland et al., 2000). The orthogneisses are interlayered
663 with comparatively heterogeneous layered gneisses (referred to as banded gneiss in the field),
664 commonly migmatitic, and generally fine-grained. The layered gneisses derive probably from
665 both volcanic and sedimentary protoliths. Thin sulfide-rich or graphite-rich schistose layers
666 are common (falhbands) (Broekmans et al., 1994; Gammon, 1966). The protoliths of the
667 orthogneisses range in age from 1575 ± 44 to 1460 ± 21 Ma, with two frequency maxima
668 around 1545 and 1505 Ma (Fig. 7) (Andersen et al., 2004a; Bingen and Viola, 2018; Engvik
669 et al., 2016). The orthogneisses have tholeiitic to low-K calc-alkaline geochemical signature,
670 typical of volcanic arc magmatism (Andersen et al., 2004a). Their Hf isotopic signature is
671 very radiogenic, with an average εHf = +8.8 (+7 < εHf < +11), approaching the depleted
672 mantle reservoir at 1550 Ma (Fig. 8) (Andersen et al., 2002b). (ii) The metasedimentary
673 complexes (Fig. 6) consist of coarse quartzite, interlayered with mica schist, sillimanite gneiss
674 and sulfide-rich schist (Morton, 1971; Nijland et al., 2014; Nijland et al., 1993). They host
675 metasomatic rocks such as orthoamphibole-cordierite gneiss, talc schist, albitite, scapolitite
676 and dolomite, generally located at the interface with gabbro bodies (Dahlgren et al., 1993;
677 Engvik et al., 2014; Munz, 1990; Munz et al., 1994). Deposition of the sedimentary protoliths
678 took place after 1467 ± 33 Ma (detrital zircon U–Pb data in quartzite samples), implying that
679 they represent part of a cover to the orthogneiss basement (Åhäll et al., 1998; Bingen et al.,
680 2001).

681 **3.4.2 Pre- to early-Sveconorwegian plutonism**

682 The orthogneiss and quartzite-rich metasedimentary complexes are intruded by variably sized,
683 gabbroic plutons. These plutons are commonly zoned, with (sub)ophitic picritic gabbro in the
684 core and garnet amphibolite along the margin (Munz and Morvik, 1991). Two such gabbro
685 plutons have been dated by the Sm-Nd method at 1224 ± 15 and 1207 ± 14 Ma (Morud and
686 Vestre Dale gabbro, not represented in Fig. 7; deHaas et al., 2002b; Munz and Morvik,
687 1991), and two have been dated with the U–Pb method at 1164 ± 12 and 1149 ± 7 Ma
688 (Vinoren and Ringsjø; Bingen and Viola, 2018; Engvik et al., 2011). Felsic intrusive rocks are
689 quite common. They include thin gneissic units ranging in age from 1178 ± 9 to 1149 ± 8 Ma
690 (Andersen et al., 2004b; Bingen and Viola, 2018; Engvik et al., 2016), and also, in Bamble,
691 larger plutons ranging in age from 1152 ± 11 to 1140 ± 13 Ma (Gjeving, Ubergsmoen,
692 Hovdefjell-Vegårshei plutons, Fig. 6) (Bingen and Viola, 2018). These metaplutons are
693 characterized by a weakly foliated magmatic charnockite facies in the centre and a garnet-
694 bearing augen gneiss facies at the margin (Touret, 1971a, b), and therefore place a maximum
695 age bracket for the high-grade deformation and metamorphism in Bamble.

696 **3.4.3 Sveconorwegian orogeny**

697 The Sveconorwegian overprint in the Bamble and Kongsberg lithotectonic units is typified by
698 a steep to subvertical foliation, isoclinal and highly transposed folds and a penetrative tectonic
699 layering (Bingen and Viola, 2018; Slagstad et al., 2020; Starmer, 1985, 1991). These features
700 are interpreted as evidence for roughly orthogonal, syn-metamorphic shortening, oriented E-
701 W for Kongsberg and NW–SE for Bamble (Bingen and Viola, 2018). A steep stretching
lineation on the steep foliation planes suggests a component of near-vertical stretching. Inside the Kongsberg lithotectonic unit, the N–S trending Hokksund-Solumsmo shear zone (Starmer, 1985) is characterized by a component of sinistral strike-slip shearing that overprinted and thus postdates the orthogonal shortening (Scheiber et al., 2015).

Metamorphic grade increases across strike, northeastwards in Kongsberg and southeastwards in Bamble. In Kongsberg, it increases from epidote-amphibolite facies to upper amphibolite-facies conditions, with local occurrences of granulite-facies rocks towards the northeast (Fig. 6). In Bamble, the grade increases from amphibolite-facies to granulite-facies conditions towards the southeast, i.e. towards the coast (Tromøy and Hisøy islands; Fig. 6) (Clough and Field, 1980; Harlov, 2000; Knudsen, 1996; Nijland et al., 2014; Nijland and Maijer, 1993; Touret, 1971a). However, patches of granulite facies rocks are scattered throughout the amphibolite-facies domain of Bamble (Mosleiken granulite; Fig. 6), underscoring the importance of fluid activity on mineral parageneses (Engvik et al., 2016; Nijland et al., 1998). The granulite-facies rocks record peak pressure-temperature values of 1.15 GPa and 850 °C, followed by hydration and decompression to 0.8 GPa – 740 °C (Engvik et al., 2016).

Zircon and monazite U–Pb data constrain the peak of amphibolite- and granulite-facies metamorphism between 1147 ±12 and 1122 ± 8 Ma in both the Bamble and Kongsberg lithotectonic units (Bingen et al., 2008b; Bingen and Viola, 2018; Cosca et al., 1998; Engvik et al., 2016; Knudsen et al., 1997). In coastal Bamble, the granulite-facies Tromøy complex (Fig. 6) consists of low-K calc-alkaline enderbitic gneisses depleted in large ion-lithophile elements (LILE) (Cooper and Field, 1977; Field et al., 1980; Knudsen and Andersen, 1999). Zircon U–Pb data demonstrate that the protoliths formed between 1575 ± 44 and 1544 ± 14 Ma while the granulite facies overprint took place between 1147 ± 12 and 1132 ± 6 Ma (Bingen and Viola, 2018). These data show that the volcanic arc magmatism belongs to the
727 Gothian orogenic cycle, while the granulite-facies metamorphism is early-Sveconorwegian in
728 age (Andersen et al., 2004a; Bingen et al., 2008c; Field et al., 1985).

729 Titanite U-Pb dates and a trail of monazite dates in gneisses range from 1107 ± 9 to 1091
730 ± 2 Ma (Bingen et al., 2008b; Cosca et al., 1998; deHaas et al., 2002a) while hornblende
731 <sup>40</sup>Ar/<sup>39</sup>Ar plateau ages range from 1099 ± 3 to 1079 ± 5 Ma (Cosca et al., 1998; Cosca and
732 O'Nions, 1994). These dates are related to regional cooling and exhumation. Monazite, titanite
733 and rutile in albitite record at least two phases of fluid-rock interaction below 550 °C
734 (metasomatism), between 1104 ± 5 and 1078 ± 3 Ma (Engvik et al., 2017; Engvik et al., 2011;
735 Munz et al., 1994), while gadolinite-columbite data in pegmatite record intrusion of a small
736 batches of fluid-rich melt between 1094 ± 11 and 1082 ± 5 Ma (Müller et al., 2017; Scherer et
737 al., 2001). These data imply regional scale fluid mobility after the peak of metamorphism and
738 deformation.

739 Rare lamprophyre dykes with near vertical attitude and non-foliated chilled margins
740 crosscut at high angle the regional foliation of the host gneiss. One such dyke yields an
741 intrusion age of 1033 ± 12 Ma and thus provides both a minimum bracket for the steep, high-
742 grade fabric of the host gneiss and the age of a batch of ultrapotassic mafic magmatism
743 (Bingen and Viola, 2018). The large non-foliated Herefoss granite pluton formed at 920 +16/-
744 27 Ma (Fig. 5) (Andersen et al., 2002a).

### 3.4.4 Kongsberg–Idefjorden boundary zone

746 The Kongsberg–Idefjorden boundary zone is marked by a c. 500 m thick amphibolite-facies
747 shear zone made of banded gneiss of mafic composition, characterized by steeply dipping
748 foliation bearing a moderately to steeply plunging lineation (Bingen and Viola, 2018). It
749 follows the lithological contact between metagreywackes of the Veme Complex and
750 orthogneisses of the Kongsberg Complex (Viola et al., 2016).
751 **3.4.5 Kongsberg-Telemarkia boundary zone**

752 The Sokna-Saggrenda Shear Zone (Fig. 2) (Starmer, 1985) is a N–S trending, east-dipping, up
753 to 2 km thick multiphase shear zone. It is largely hosted within and along the eastern margin
754 of a > 100 km long belt of foliated granite, dated between 1170 ± 11 and 1146 ± 5 Ma (Fig. 6)
755 (Scheiber et al., 2015). This granite constitutes the footwall of the shear zone and is part the
756 Telemarkia lithotectonic unit. Three post-1170 Ma ductile deformation phases have been
757 identified in the shear zone (Scheiber et al., 2015). (i) The earliest structures accommodate
758 top-to-the-west kinematics and relate to thrusting of Kongsberg over Telemarkia. (ii) These
759 are selectively reactivated in a sinistral fashion along mylonitic to ultramylonitic shear zones.
760 The sinistral shear zones possibly record the same deformation as the N–S trending, steeply
761 dipping Hokksund-Solumsmo mylonite zones inside the Kongsberg lithotectonic unit,
762 showing evidence for sinistral transpressive shearing. (iii) Extensional top-to-the-east sense of
763 shear. A brittle zone overprinting this long-lived ductile deformation zone and traditionally
764 referred to as the "Great Friction Breccia" (Starmer, 1985) probably represents a normal fault
765 of Permian age (Larsen et al., 2008; Scheiber et al., 2015).

766 **3.4.6 Bamble-Telemarkia boundary zone**

767 The Kristiansand–Porsgrunn Shear Zone (Fig. 2) is a c. 1–2 km thick ductile to brittle shear
768 zone juxtaposing the Bamble and Telemarkia lithotectonic units. It dips moderately to the
769 southeast and is possibly connected with an offset of the Moho under the Skagerrak sea
770 (Andersson et al., 1996). The shear zone is interpreted as a top-to-the-northwest thrust, later
771 reactivated coaxially as an extensional shear zone (Henderson and Ihlen, 2004; Mulch et al.,
772 2005; Starmer, 1991). Upper greenschist- to amphibolite-facies thrust-related structures are
773 invariably northwest vergent. These structures are associated with tabular pegmatite bodies
774 (Henderson and Ihlen, 2004). The shear zone overprints the 1132 ± 3 Ma Morkheia
775 monzonite suite exposed in the Telemarkia footwall (Heaman and Smalley, 1994; Milne and
776

Starmer, 1982), and the 1140 ± 13 Ma Hovdefjell-Vegårshei metapluton exposed on the

777

Bamble hangingwall (Bingen and Viola, 2018; Touret, 1987), implying that thrusting is

778

younger than 1132 ± 3 Ma. Extension was accommodated by thin greenschist-facies shear

779

zones with syn-kinematic muscovite porphyroblasts constraining the top-to-the-southeast

780

deformation between 891 ± 3 and 880 ± 3 Ma ($^{40}$Ar/$^{39}$Ar data) (Mulch et al., 2005). The

781

contrast in titanite U–Pb ages between the Telemarkia footwall (c. 913 to 901 Ma) and

782

Bamble hanging wall (c. 1107 to 1091 Ma) (Bingen et al., 1998; Cosca et al., 1998; deHaas et

783

al., 2002a; Heaman and Smalley, 1994) underscores the importance of normal movement

784

along the shear zone. A narrow, fully brittle, Permian, normal fault zone locally reactivates

785

the Sveconorwegian ductile precursors.

786

<figure><img src="image_25.png" /></figure>
**Figure 11.** Generalized stratigraphic columns for the Telemark supracrustal rocks in central Telemark and supracrustal rocks in the Sæsvatn-Valldal area. These columns follow the archetypal subdivision into the Rjukan, Seljord and Bandak successions (Dons, 1960; Dons and Jorde, 1978; Sigmond, 1978), and integrate results of later mapping. Main sources of stratigraphic and geochronological data: Bingen et al. (2002), Bingen et al. (2003), Corfu and Laajoki (2008), Dons (1960; 1978), Laajoki et al. (2002), Laajoki and Corfu (2007), Köykkä and Lammingen (2011), Lamminen and Köykkä (2010), Lamminen (2011), Nordgulen (1999), Sigmond (1975, 1978, 1998), and Spencer et al. (2014).

796

<figure><img src="image_26.png" /></figure>

**Figure 12.** Kernel density estimators of detrital zircon ages in metasediments of the Telemarkia lithotectonic unit compared with paragneisses and metasediments in Quebecia, Oaxaquia, and inliers in the Andes of Colombia. The Seljord succession in Telemarkia and
801 the Port au Quilles formation in the Escumins supracrustals record marine peri-Baltica and
802 peri-Laurentia sedimentation, respectively, after the 1520-1480 Ma continental generation.

803 The main peak reflects sourcing in the juvenile c. 1520-1480 Ma volcanic arcs, while the
804 diversity of older detrital zircons reflects sourcing from continental sources. The Bandak
805 succession in Telemarkia deposited after c. 1180 Ma (Eidsborg Formation after c. 1100 Ma)

806 and involved important recycling of the Seljord succession and younger magmatic rocks in
807 continental intramontane environment. Contrasting with this situation, the Oaxaquia

808 lithotectonic unit and the Inliers in the Andes of Colombia are interpreted as outboard

809 volcanic arcs formed in the ocean between Laurentia, Amazonia and Baltica after c. 1460 Ma,
810 and isolated almost entirely from continental sediment sources older than c. 1500 Ma. The
811 plots are generated with "DensityPlotter" by (Vermeesch, 2012) with a bandwidth of 10 Ma.

812 Data sources: Telemarkia, Bandak and Seljord successions: Bingen et al. (2001), de Haas et
813 al. (1999), Lamminen (2011) and Spencer et al. (2014); Quebecia: Escumins supracrustal

814 rocks, Port aux Quilles formation, Groulier et al. (2018b); Oaxaquia: granulite-facies

815 paragneisses, Solari et al. (2014); inliers in the Andes of Colombia: paragneisses, Cardona et
816 al. (2010) and Ibanez-Mejia et al. (2011).

817 -----------------
<figure><img src="image_27.png" /><figcaption>**Figure 13.** Neodymium isotopic composition of rock suites in the Telemarkia lithotectonic unit, expressed as εNd (initial value) as a function of time. Each symbol represents one sample. Magmatic rocks are represented at their probable time of crystallization and metasedimentary rocks at their probable time of deposition (to improve legibility, each symbol is assigned a random scatter lower than ± 8 Ma along the time axis). Interpretations of the distribution of data are discussed in the text. (a) Rock suites in the Telemarkia lithotectonic unit. (b) Comparison between Telemarkia, Quebecia (Grenville orogen, Canada), the Composite Arc and Frontenac-Adirondack belt (Grenville orogen, Canada, USA), the Oaxaquia lithotectonic units (Mexico), and the inliers in the Andes of Colombia. These five lithotectonic units have</figcaption></figure>

818

819

820

821

822

823

824

825

826

827
similar crustal evolution vectors. Sources of data: Hunnedalen dolerites at c. 870 Ma: (Maijer and Verschure, 1998); Rogaland AMC suite at 930 Ma: (Barling et al., 2000; Bolle et al., 2003a; Demaiffe et al., 1986; Menuge, 1988; Nielsen et al., 1996; Robins et al., 1997); Schiellerup et al., 2000); high-alumina orthopyroxene megacrysts (HAOM) in anorthosite plutons, 1040-930 Ma: (Bybee et al., 2014; Demaiffe et al., 1986); HBG granitoids, 985-925 Ma: (Andersen et al., 2001; Bogaerts et al., 2003; Demaiffe et al., 1990; Menuge, 1985, 1988; Vander Auwera et al., 2003; Vander Auwera et al., 2014a); Feda suite at 1050 Ma: (Bingen et al., 1993; Menuge, 1988; Vander Auwera et al., 2011); other magmatic rocks in Telemarkia: (Andersen et al., 2001; Brewer et al., 2002; Brewer et al., 2004; Brewer and Menuge, 1998; Menuge, 1985, 1988; Vander Auwera et al., 2003); metasedimentary rocks: (Andersen and Laajoki, 2003; deHaas et al., 1999); Oaxaquia: (Lawlor et al., 1999; Ruiz et al., 1988; Weber and Köhler, 1999); Inliers in the Andes of Colombia: (Cordani et al., 2005; Ibanez-Mejia et al., 2015); Quebecia: (Dickin, 2000; Dickin and Higgins, 1992; Groulier et al., 2018a; Groulier et al., 2018b); Composite Arc and Frontenac-Adirondack belt: (Chiarenzelli et al., 2010; Daly and McLelland, 1991; Dickin et al., 2010; Marcantonio et al., 1990; McLelland et al., 1993; Valentino et al., 2019).

844 ------------------
<figure><img src="image_28.png" /></figure>

**Figure 14.** Comparison of the geochemical signature between three diagnostic magmatic suites intruded between 1066 and 916 Ma in the Telemarkia lithotectonic unit: c. 1050 Ma high-K calc-alkalic Feda plutonic suite of the Sirdal magmatic belt and 1030 Ma Fennefoss pluton (Bingen, 1989; Pedersen, 1981; Vander Auwera et al., 2011), 986–926 Ma hornblende-biotite granite (HBG) suite with ferro-potassic calc-alkalic to alkali-calcic signature (Bogaerts et al., 2003; Vander Auwera et al., 2003; Vander Auwera et al., 2014a) and the 937–916 Ma anorthosite-mangerite-charnockite (AMC) suite with ferro-potassic alkalic signature (Bolle and Duchesne, 2007; Charlier et al., 2010; Duchesne and Wilmart, 1997; Vander Auwera et al., 2014a; Vander Auwera et al., 1998; Wilmart et al., 1989). For the AMC suite, 3 different trends are recognized based on the mineralogy (two-pyroxene and fayalite trends) or their belonging to a specific intrusion (apophysis of the Bjerkreim-Sokndal intrusion).

**3.5 Telemarkia lithotectonic unit**

The western part of the Sveconorwegian orogen can be considered as one single lithotectonic unit, c. 230 x 300 km long, named Telemarkia (Fig. 2) (Bingen et al., 2005). The Telemarkia lithotectonic unit comprises low-grade supracrustal rocks preserved in several syncline structures, structurally overlying amphibolite- to granulite-facies gneiss complexes, and hosts
862 voluminous plutons (Fig. 4; Fig. 5; Fig. 6). The gneiss complexes comprise orthogneisses

863 with subordinate paragneisses. The largest and most complete tract of supracrustal rocks,

864 called the Telemark supracrustal rocks, is exposed in a 60 km wide area in central Telemark

865 (Fig. 4; Fig. 6). Original mapping showed that stratigraphic relationships and deposition

866 structures are well preserved in the Telemark supracrustal rocks, and defined three groups or

867 successions separated by unconformities, which are, from bottom to top, the Rjukan, Seljord

868 and Bandak successions (Fig. 11, Fig. 12) (Dons, 1960; Dons and Jorde, 1978; Sigmond,

869 1978). Other supracrustal sequences are described in several other syncline structures, less

870 than 25 km wide, in the Ullensvang, Sauda, Grjotdokki-Nesflaten, Sæsvatn-Valldal and

871 Nissedal areas (Fig. 4).

872 **3.5.1 Telemarkian evolution**

873 Rapid generation of juvenile continental crust is recorded by voluminous magmatism between

874 1521 ± 6 and 1476 ± 13 Ma, hosted in both the low-grade successions and in the gneiss

875 complexes (Fig. 7) (Bingen et al., 2008a; Bingen et al., 2005; Laajoki and Corfu, 2007;

876 Pedersen et al., 2009; Roberts et al., 2013). This event is called the Telemarkian accretionary

877 orogeny and it is geographically zoned.

878 In the west, in the Suldal area, gneisses and granitoids are characterized by a calc-alkaline

879 geochemical signature, with s supra-chondritic Hf isotopic signature at 1500 Ma (average εHf

880 = +5.7; Fig. 8) (Pedersen et al., 2009; Roberts et al., 2013). These are interpreted to reflect

881 volcanic arc magmatism (the Suldal arc; Roberts et al., 2013).

882 In the east, in the Telemark area, magmatism is bimodal and typified by the Rjukan

883 bimodal metavolcanic rocks at the base of the Telemark supracrustal rocks (Vemork basalt vs.

884 Tuddal rhyolite dated between 1512 ± 10 and 1495 ± 2 Ma, Fig. 11) and coeval plutonism

885 (Bingen et al., 2005; Laajoki and Corfu, 2007). This magmatism is characterized by a within-

886 plate geochemical signature and moderately supra-chondritic Nd isotopic signatures (+1.1 <
887 $\epsilon_{Nd}(1500\ \text{Ma}) < +4.3$) (Fig. 13). It is interpreted to reflect back-arc rifting (the Rjukan rift basin),
888 continentwards of the active arc (Brewer and Menuge, 1998; Köykkä and Lamminen, 2011;
889 Lamminen and Köykkä, 2010; Roberts et al., 2013).

890 The Telemarkian orogenic event cannot be demonstrated to be associated with high-grade
891 metamorphism. The Seljord succession overlying the Rjukan succession is a c. 8 km thick,
892 shallow marine sedimentary succession, dominated by quartzite (Fig. 11, Fig. 12) (Köykkä
893 and Lamminen, 2011). The Seljord succession was deposited during a transgressive cycle,
894 interpreted as reflecting thermal subsidence after magmatism, between 1410 ± 24 Ma (detrital
895 zircon U–Pb data) and 1347 ± 4 Ma (U–Pb age of intrusive dolerite dyke; Corfu and Laajoki,
896 2008; Köykkä and Lamminen, 2011; Lamminen and Köykkä, 2010).

897 *3.5.2 Pre- to early-Sveconorwegian evolution*

898 The Telemarkia lithotectonic unit hosts several generations of gneissic plutonic rocks in the
899 1280–1240, 1220–1180 and 1180–1145 Ma time intervals, with frequency maxima at c. 1280,
900 1260, 1210, 1170 and 1150 Ma (Fig. 6, Fig. 7) (Andersen et al., 2007; Bingen et al., 2003;
901 Corfu and Laajoki, 2008; Heaman and Smalley, 1994; Pedersen et al., 2009; Scheiber et al.,
902 2015).

903 In the southeast of the Telemarkia lithotectonic unit, a voluminous c. 60 x 120 km gneiss
904 complex consists of amphibolite-facies, NE–SW trending, moderately to weakly foliated
905 granitic gneiss and granitoids, named in different areas Drivheia gneiss (Heaman and
906 Smalley, 1994) and Vråvatn complex (Fig. 6) (Andersen et al., 2007). This complex is
907 dominated by c. 1220–1190 Ma plutonic rocks with a within-plate geochemical signature
908 (Andersen et al., 2007; Bingen and Viola, 2018; Heaman and Smalley, 1994) and a supra-
909 chondritic (radiogenic) Hf isotopic signature (+9 < $\epsilon_{Hf}$ < +10, in zircon from 4 samples, Fig.
910 8) approaching the depleted mantle reservoir value at 1210 Ma ($\epsilon_{Hf}$ = +12) (Andersen et al.,
911 2007).
The Sæsvatn–Valldal and Nissedal supracrustal rocks are two low-grade basalt-dominated successions (Fig. 6, Fig. 11), exposed in two c. 15 km wide syncline. Basalt is interlayered with felsic volcanic and clastic sedimentary rocks and intruded by fine-grained granite sills and dykes (Dons and Jorde, 1978; Sigmond, 1975). In the Sæsvatn–Valldal succession, the basalts are overlying rhyolites and porphyries dated to between 1275 ± 8 and 1259 ± 2 Ma, themselves unconformably overlying the 1520–1480 Ma gneissic basement (Bingen et al., 2002; Brewer et al., 2004). In the Nissedal succession, the basalts overly the 1219 ± 8 to 1202 ± 9 Ma Vråvatn complex and host fine-grained granite sheets, one of which yields an intrusion age of 1196 ± 6 Ma (Bingen and Viola, 2018). The Nissedal and Sæsvatn–Valldal successions are interpreted as near-coeval bimodal (mafic dominated) continental successions with an age close to 1210 Ma, coeval with the Drivheia and Vråvatn gneisses in the underlying gneiss complex (Andersen et al., 2007; Heaman and Smalley, 1994).

In the Telemark supracrustal rocks, the c. 3.5 km thick Bandak succession rests over both the Rjukan and Seljord successions (Køykkä, 2011; Laajoki et al., 2002), above a first order unconformity locally decorated by a regolith (Køykkä and Laajoki, 2009) (Fig. 6, Fig. 11, Fig. 12). The succession includes at least two internal unconformities, implying active tectonism during sedimentation (Laajoki, 2002; Laajoki et al., 2002). The lower part of the Bandak succession consists of bimodal volcanic rocks interlayered with sediments (Køykkä, 2011). The mafic rocks (Morgedal and Gjuve metabasalts) have a within-plate geochemical signature (Brewer et al., 2002; Spencer et al., 2014). The felsic volcanic rocks range in age from 1169 ± 9 to 1145 ± 4 Ma (Bingen et al., 2003; Laajoki et al., 2002). The upper part of the Bandak succession consists of exclusively sedimentary rocks. These are the Heddal Group, Eidsborg Formation and Kalhovd Formation, deposited after 1116 ± 24, 1103 ± 14 and 1054 ± 22 Ma respectively (Fig. 5, Fig. 11) (detrital zircon U–Pb data; Bingen et al., 2003; deHaas et al., 1999; Lamminen, 2011; Spencer et al., 2014).
The sedimentary rocks of the entire Bandak succession are generally immature, coarse-grained to conglomeratic, and of limited lateral extent. They are interpreted as alluvial fan-, braided fluvial- and locally eolian deposits, accumulated in continental fault-bounded intermontane extensional basins (Bingen et al., 2003; Köykkä, 2011; Lamminen, 2011; Spencer et al., 2014). Syn-sedimentary normal faults are well documented (Lamminen, 2011).

### 3.5.3 Sveconorwegian magmatism

After 70 Myr of quiescence, magmatism resumed at c. 1065 Ma with formation of the c. 50 km wide – 170 km long, orogen-parallel, NNW-SSE trending, Sirdal magmatic belt in the Agder area (Fig. 5) (Bingen et al., 2015; Coint et al., 2015; Granseth et al., 2020; Slagstad et al., 2013). This belt is a composite granitoid batholith, comprising mainly elongate and variably foliated plutons of granodiorite, granite and leucogranite. Slivers of heterogeneous gneiss interleaved within granitoid plutons are interpreted as xenoliths or panels of wall-rocks (Coint et al., 2015). The granitoids intruded under pressure conditions of 0.38–0.48 GPa (Coint et al., 2015) between 1066 ± 10 and 1020 ± 15 Ma (Bingen et al., 2015; Bingen and van Breemen, 1998a; Coint et al., 2015; Möller et al., 2002; Slagstad et al., 2018; Slagstad et al., 2013). A large portion of the belt comprises silica-rich biotite granite and leucogranite.
Foliated plutons of biotite + amphibole K-feldspar-phyric quartz-monzonite to granodiorite are specifically called the Feda suite (1050 ± 8 Ma) and Fennefoss augen gneiss (1031 ± 2 Ma) (Fig. 5) (Bingen and van Breemen, 1998a). These are characterized by a magnesian, high-K, high-Sr-Ba, calc-alkaline geochemical signature and locally host ultrapotassic (lamprophyre) mafic layers and enclaves (Fig. 14) (Bingen et al., 1993; Bingen and van Breemen, 1998a).

After 985 Ma, large plutons with a distinctly ferroan geochemical signature were emplaced (Fig. 5, Fig. 14) (Andersen et al., 2001; Granseth et al., 2020; Vander Auwera et al., 2011). These plutons are weakly- to non-foliated, have sharp contacts to their wall-rock and are well
defined on aeromagnetic maps by positive anomalies (Slagstad et al., 2018). Two main ferroan suites are defined: a ferro-potassic hornblende-biotite-granitoid (HBG) suite and an orthopyroxene-bearing anorthosite-mangerite-charnockite (AMC) suite (Fig. 5, Fig. 14) (Bogaerts et al., 2003; Duchesne and Wilmart, 1997; Vander Auwera et al., 2003; Vander Auwera et al., 2011; Vander Auwera et al., 2014a). The HBG suite formed between 986 ± 2 and 926 ± 4 Ma and is exposed in the area of the Sirdal magmatic belt and eastwards (Andersen et al., 2001; Andersen et al., 2007; Granseth et al., 2020; Jensen and Corfu, 2016; Sigmond, 1985; Slagstad et al., 2018; Vander Auwera et al., 2011; Vander Auwera et al., 2014a). The AMC suite formed between 937 ± 1 and 916 ± 9 Ma and is restricted to the southwestern end of the Telemarkia lithotectonic unit (Fig. 5) (Bolle et al., 2018; Schärer et al., 1996; Vander Auwera et al., 2011; Vander Auwera et al., 2014a). A few plutons (Farsund and Kleivan plutons, 931 ± 2 and 936 ± 1 Ma) are composite HBG-AMC plutons, with charnockitic and non-charnockitic facies, reflecting tapping of distinct sources into one pluton (Vander Auwera et al., 2014a).

The Rogaland AMC suite (Fig. 5) consists of three large anorthosite plutons (Egersund-Ogna, Håland-Helleren, Åna-Sira anorthosites), two satellite leuconorite plutons (Hidra and Garsaknatt leuconorites), a layered intrusion (Bjerkreim-Sokndal layered intrusion), and volumetrically minor sills and dykes of jotunite and ilmenite-norite, all emplaced during a short lived magmatic event between c. 932 and 916 Ma (Charlier et al., 2006; Duchesne et al., 1985; Duchesne et al., 1989; Schärer et al., 1996; Vander Auwera et al., 2011).

The Egersund-Ogna anorthosite pluton exhibits an isotropic core and a foliated margin, characterized by a syn-magmatic fabric parallel to the contact. The centre of the pluton is made up of anorthosite and leuconorite with a granulated matrix of plagioclase (An40–An50), hosting 1–3 m large aggregates of plagioclase (up to An55) and high-alumina orthopyroxene megacrysts (HAOM, En75) (Charlier et al., 2010). The high aluminium and chromium
contents (up to 8.5 wt% Al₂O₃ and 1500 ppm Cr) of the orthopyroxene megacrysts indicate a pressure of crystallization of c. 1.1 GPa for the megacrysts, contrasting with the ambient pressure of 0.5 GPa for the matrix minerals (2-3 wt% Al₂O₃ in matrix orthopyroxene). The anorthosite plutons intruded as a plagioclase-dominated crystal mush lubricated by melt, from the base of the crust (1.1 GPa) to the middle of the crust (0.5 GPa) (Barnichon et al., 1999; Charlier et al., 2010; Duchesne et al., 1999). The orthopyroxene megacrysts with the highest aluminum content (> 8 wt% Al₂O₃) define a Sm-Nd isochron with an age of 1041 ± 17 Ma (Bybee et al., 2014), pointing either to inheritance (Vander Auwera et al., 2014b) or protracted ponding of mafic magma at the base of the crust (Bybee et al., 2014).

The Bjerkreim-Sokndal layered intrusion (931 ± 7 Ma) can be subdivided into a layered lower part and a non-layered upper part. The lower part comprises five macrocyclic units of cumulates (Barling et al., 2000; Duchesne, 1972; Nielsen et al., 1996; Robins et al., 1997).

The upper part comprises, fractionated and wall-rock-contaminated, mangerite and charnockite (Duchesne and Wilmart, 1997; Nielsen et al., 1996). The Bjerkreim-Sokndal intrusion intruded at pressure conditions of ≤ 0.5 GPa (Vander Auwera and Longhi, 1994). It forms a syncline (lopolith), the formation of which is attributed to gravity-driven subsidence of the central part of the intrusion (Bolle et al., 2000; Bolle et al., 2002; Paludan et al., 1994).

Undeformed pegmatites intruded between c. 914 and 900 Ma. They include the Evje-Iveland rare-mineral pegmatite field (Müller et al., 2017; Pasteels et al., 1979; Scherer et al., 2001; Seydoux-Guillaume et al., 2012).

### 3.5.4 Sveconorwegian metamorphism

As outlined above, the supracrustal rocks in the centre of the Telemarkia lithotectonic unit were affected by greenschist to epidote-amphibolite facies metamorphism and deformed by open to tight folding. Basalt in the Sæsvatn-Valldal succession (Fig. 6) was deformed under
1011 epidote-amphibolite facies conditions at c. 1032 ± 2 Ma and faulted at 1017 ± 2 Ma
1012 (molybdenite Re-Os data; Stein and Bingen, 2002).

1013 In the gneiss complexes, the metamorphic grade typically reached upper amphibolite-
1014 facies conditions, with widespread migmatitization between 1026 ± 14 and 1005 ± 7 Ma
1015 (zircon and monazite U–Pb data; Bingen et al., 2008b; Coint et al., 2015). Granitoids of the
1016 Sirdal magmatic belt (c. 1065–1020 Ma) are commonly moderately deformed (Coint et al.,
1017 2015). Locally, they contain zircons with rims recording a hydrothermal to metamorphic
1018 overprint at c. 1016 Ma (Knaben Mo district ; Bingen et al., 2015).

1019 The metamorphic grade increases southwestwards towards the Rogaland AMC complex,
1020 structurally downwards, across to the N–S trending and E-dipping regional fabric (Fig. 5)
1021 (Bingen and van Breemen, 1998b; Maijer, 1987; Slagstad et al., 2018; Tobi et al., 1985).
1022 Metamorphism was coeval with the formation of lithological banding, tight to isoclinal
1023 folding and migmatitization. Two concentric granulite facies zones are defined: the
1024 orthopyroxene zone and the osumilite zone close to the AMC complex (Fig. 5). Osumilite is
1025 diagnostic of water poor, low-pressure, ultrahigh temperature (UHT; T > 900 °C) granulite-
1026 facies conditions (Harley, 2008; Holland et al., 1996).

1027 Zircon and monazite U–Pb geochronology from a diversity of granulite-facies samples
1028 gave apparent ages spreading between c. 1045 and 900 Ma (Bingen et al., 2008b; Bingen and
1029 van Breemen, 1998b; Laurent et al., 2018a; Möller et al., 2002, 2003; Slagstad et al., 2018;
1030 Tomkins et al., 2005). Insight into the pressure-temperature-time evolution of this protracted
1031 metamorphism requires careful linkage of petrography, phase equilibrium modelling,
1032 geochronology and trace-element characterization of zircon and monazite. Typical samples
1033 inside the orthopyroxene zone reached peak conditions of 0.5 GPa – 880 °C between c. 1040
1034 and 1010 Ma (Laurent et al., 2018b). Rims of neocrystallized zircon in such samples spread
1035 from 1045 to 955 Ma, supporting 90 Myr of melt-present conditions (Laurent et al., 2018a).
1036 In the osumilite zone, the onset of migmatitization, associated with biotite and sulfide mineral breakdown, is recorded by sulfate-rich monazite cores in an osumilite-bearing paragneiss at 1037
1038 1034 ± 6 Ma (Laurent et al., 2016). In a (quartz- and garnet-free) sapphirine + orthopyroxene sample (Fig. 9; Ivesdal locality), a Y-rich monazite (5–7 wt% Y₂O₃) further constrains 1039
1040 temperature higher than 900°C between 1029 ± 9 and 1006 ± 8 Ma (Laurent et al., 2018b), in 1041 accordance with zircon data (1010 ± 7 to 1006 ± 4 Ma) (Drüppel et al., 2013). The breakdown 1042 of the peak sapphirine + orthopyroxene assemblage into a cordierite + hercynite assemblage 1043 implies a clockwise P-T path with a decompression between 0.6 GPa – 920 °C and 4.5 GPa – 1044
1045 900°C (Fig. 9) (Blereau et al., 2017; Laurent et al., 2018b). This decompression is best captured by a garnet-bearing sample from the osumilite zone that contains Y-rich monazite 1046
1047 recording garnet breakdown into cordierite + hercynite + orthopyroxene, pinning a robust P-T-t point at 0.4 GPa – 910 °C – 930 ± 6 Ma (Laurent et al., 2018b). Together, the data give 1048
1049 evidence for two events of low-pressure granulite-facies metamorphism peaking at UHT conditions, the first event (M1) between c. 1030 and 1005 Ma, and the second (M2) at c. 930 1050
1051 Ma, associated with formation of osumilite (Blereau et al., 2017; Drüppel et al., 2013; Laurent et al., 2018b; Laurent et al., 2016). Minor exhumation (c. 6 km) took place between the two.
1052

1052 These two events were penecontemporaneous with magmatic activity (Laurent et al., 2018b; Slagstad et al., 2018). The first M1 event started with dehydration melting (c. 1034 Ma) coeval with intrusion of the Sirdal magmatic belt (c. 1065-1020 Ma) and associated underplating (c. 1040 Ma), and peaked at the end and after this magmatic event (1030–1005 Ma). The second M2 event (c. 930 Ma) was coeval with intrusion of the AMC suite. This correlation strongly suggests that magmatism and metamorphism had a common heat source in the mantle. The lag between magmatism and peak metamorphism for M1 may reflect temperature buffering by melt until melt migration effectively took place.
1053
1054
1055
1056
1057
1058
1059
1060 The M2 metamorphic event was followed by regional scale cooling, dated by titanite U-Pb
1061 data at 918 ± 2 Ma (Bingen and van Breemen, 1998b). Amphibole <sup>40</sup>Ar/<sup>39</sup>Ar apparent ages
1062 scatter between 1059 ± 8 and 853 ± 3 Ma (Bingen et al., 1998). The main cluster at 871 ± 10
1063 Ma overlaps with biotite Rb-Sr ages (Verschure et al., 1980) and is interpreted as a cooling
1064 age.

1065 *3.5.5 The Mandal-Ustaoset fault and shear zone*

1066 The Mandal-Ustaoset fault and shear zone is a N–S trending structure inside the Telemarkia
1067 lithotectonic unit (Fig. 2). It includes a precursor ductile shear zone and a set of later brittle
1068 normal faults (Sigmond, 1985). In its northern segment, it is an east dipping (c. 45°) normal
1069 (extensional) shear zone, juxtaposing the amphibolite-facies Hardangervidda gneiss complex
1070 in the west against the low-grade intramontane basin hosting the Kalhovd Formation (≤ 1054
1071 ± 22 Ma) in the east (Sigmond and Ragnhildstveit, 2004). Towards the south, the Mandal-
1072 Ustaoset fault and shear zone merges into an amphibolite-facies N–S trending banded gneiss
1073 unit on the eastern side of an elongate pluton of the Feda suite (1049 ± 8 Ma, Mandal augen
1074 gneiss; Bingen and van Breemen, 1998a). The Mandal-Ustaoset fault and shear zone still
1075 requires detailed kinematic and geochronological characterization.

1076 **4 Discussion**

1077 **4.1 U-Pb and Lu-Hf evidence for continental growth at the margin of Fennoscandia**

1078 The continental crust exposed in the Sveconorwegian orogen was formed after 1900 Ma
1079 (Åhäll and Connelly, 2008; Andersen et al., 2004a; Bingen et al., 2005; Bingen and Viola,
1080 2018; Petersson et al., 2015b; Roberts and Slagstad, 2015; Roberts et al., 2013). The age of
1081 the dominant magmatic suites in the different lithotectonic units decreases towards the west
1082 (Fig. 7). The oldest major magmatic suites in each lithotectonic unit are dated between 1710
1083 and 1660 Ma in the Eastern Segment, 1660 and 1520 Ma in the Idefjorden lithotectonic unit,
1575 and 1480 Ma in the Bamble and Kongsberg lithotectonic units, and 1520 and 1480 Ma in the Telemarkia lithotectonic unit. The Lu-Hf isotopic signature of igneous zircon in these magmatic suites (Fig. 8) becomes more radiogenic (more positive εHf values) westward in the orogenic belt, with average initial εHf values increasing from +3.0 in the Eastern Segment (1700 Ma) to +8.8 in the Bamble-Kongsberg lithotectonic units (1550 Ma), and back to +5.7 in the Telemarkia lithotectonic unit (1500 Ma; Fig. 8) (Andersen et al., 2002b; Pedersen et al., 2009; Petersson et al., 2015a; Petersson et al., 2015b; Roberts et al., 2013). The geochemical signature of these different magmatic suites generally ranges from calc-alkalic to alkali-calcic (references above), suggesting that the continental lithosphere was generated dominantly in a supra-subduction (accretionary) geodynamic setting between 1710 and 1480 Ma (Åhäll and Connelly, 2008; Andersen et al., 2004a; Petersson et al., 2015a; Petersson et al., 2015b; Roberts et al., 2013).

The weakly positive initial εHf values in the Eastern Segment (Fig. 8) imply significant recycling of older Paleoproterozoic (Svecokarelian) continental crust in the genesis of the 1710–1660 Ma magmatic suites (Petersson et al., 2015a). The more positive initial εHf values westwards imply, instead, that the four lithotectonic units to the west of the Mylonite Zone were generated in more juvenile volcanic arc and back arc environment, away from old Paleoproterozoic continental lithosphere (Andersen et al., 2002b; Petersson et al., 2015b; Roberts et al., 2013). The variability within and between these units can be accounted for by a change from an advancing to a retreating subduction system or, alternatively, a variable contribution of metasedimentary components incorporated in the subduction system along the oceanic lower plate (Andersen et al., 2002b; Petersson et al., 2015b; Roberts et al., 2013).

The age and isotopic trends of magmatism in the 1710–1480 Ma interval (Fig. 7, Fig. 8) are compatible with incremental westward growth of the continental lithosphere at the margin
1108 of Fennoscandia. This is compatible with any orogenic model interpreting the lithotectonic

1109 units as endemic to the margin of Fennoscandia (Fig. 3 b, d, e, f).

1110

<figure><img src="image_29.png" /></figure>

High-K calc-alkaline suites

Massif Central, France
Hercynian, 335-300 Ma

Caledonides, Scotland-Ireland
Caledonian, 430-400 Ma

Tismana pluton, Romania, 570 Ma

Touareg shield, Mali-Niger
Pan-African, 700-580 Ma

Sveconorwegian Sirdal Magmatic Belt

Fennefoss Am-Bt pluton, 1030 Ma

Feda Am-Bt plutons, 1050 Ma

Sirdal Bt plutons, 1065-1020 Ma

Felsic
(n)

◆ (224)

△ (95)

■ (18)

◎ (183)

Mafic
(n)

◆ (279)

△ (52)

■ (13)

◆ (9)

◎ (65)

◎ (53)

● (7)
1112 **Figure 15.** Comparison of the geochemical signature of the Sirdal magmatic belt (c. 1065–1020 Ma) in the Telemarkia lithotectonic unit, and syn-collisional high-K calc-alkaline magmatic suites in younger collision orogenic belts. Large symbols represent granitoids while small symbols represent associated minor mafic sills, dykes and enclaves. The Sirdal magmatic belt is divided into the magnesian amphibole-biotite Feda plutonic suite hosting minor volume of ultrapotassic enclaves, the ferroan amphibole-biotite-bearing Fennefoss pluton and more silica-rich biotite-bearing foliated plutons (Bingen, 1989; Pedersen, 1981; Slagstad et al., 2013; Vander Auwera et al., 2011). For comparison, the Hercynian high-K calc-alkaline plutons of the Massif Central in France associated with "vaugnerites" (compilation: Moyen et al., 2017), the Caledonian high-K, high Ba-Sr plutons of Scotland and Ireland associated with "appinites" (Clemens et al., 2009; Ghani and Atherton, 2006), the Neoproterozoic shoshonitic Tismana pluton in the Carpathians of Romania (Duchesne et al., 1998), and the Neoproterozoic high-K calc-alkalines suites of the Touareg Shield in Mali and Niger (Liégeois et al., 1998). The figure shows a broad overlap of the geochemistry between these different suites. (a) SiO₂ vs. Na₂O+K₂O-CaO diagram (Frost et al., 2001). (b) SiO₂ vs. K₂O diagram (Peccerillo and Taylor, 1976). (c) SiO₂ vs. FeOₜot / (FeOₜot+MgO) (Frost et al., 2001). (d) SiO₂ vs. Sr+Ba diagram showing the high Sr+Ba signature of high-K calc-alkaline suites, including the Feda plutons and their ultrapotassic mafic enclaves. (e) B-A diagram (B = Fe+Mg+Ti, A = Al-(K+Na+2Ca) (expressed in gram-atoms of each element in 100 gr of material) (Debon and Le Fort, 1983; Villaseca et al., 1998) (h-P: highly peraluminous, m-P: moderately peraluminous, l-P: low peraluminous, f-P: felsic peraluminous).

1133 **4.2 Significance of high-K calc-alkaline granite plutonism**

1134 Tracing past subduction systems largely relies on tracing a subduction-related geochemical signature in magmatic rocks. In the Telemarkia lithotectonic unit, a significant component of the large (50 x 170 km) orogen-parallel Sirdal magmatic belt (c. 1065–1020 Ma) has a calc-
1137 alkaline geochemical signature (Bingen et al., 2015; Coint et al., 2015; Granseth et al., 2020; Slagstad et al., 2018; Slagstad et al., 2013). More specifically, the biotite + amphibole K-feldspar-phyric quartz-monzonite-granodiorite foliated plutons of the Feda suite are characterized by a high-K, high-Sr-Ba, magnesian, calc-alkalic geochemical trend (Fig. 14) (Bingen et al., 1993; Bingen and van Breemen, 1998a). They are associated with a small volume of ultrapotassic rocks. Calc-alkaline rocks are typically observed in active supra-subduction environment (Bateman and Chappell, 1979; Hervé et al., 2007; Pearce et al., 1984). However, high-K calc-alkaline suites are also typically representative of syn- to late-collision plutons and batholiths in collisional orogens (Fig. 15). They are well described in the Caledonian orogen (Bruand et al., 2014; Clemens et al., 2009; Ghani and Atherton, 2006; Neilson et al., 2009), the Hercynian orogen (Couzinié et al., 2016; Laurent et al., 2014; Laurent et al., 2017; Moyen et al., 2017) and the Pan-African orogens (Janoušek et al., 2010; Liégeois et al., 1998). These plutons are commonly associated with minor volumes of ultrapotassic rocks such as lamprophyre, appinite or vaugnerite (Fig. 15 b). The Sirdal magmatic belt and more specifically the Feda suite exhibits a complete overlap in major and trace element geochemical composition with syn- to late-collision high-K calc-alkaline plutons in collisional orogens (Fig. 15). Therefore, the belt could be reasonably interpreted as well as the product of syn- to late-collision magmatism.

1155 To sum up, the geochemical signature of the Sirdal magmatic belt is not fully diagnostic of a geodynamic environment. There are two alternatives. (i) It records supra-subduction magmatism as part of an active subduction system in the 1065–1020 Ma time interval. This subduction was either dipping eastwards in the context of the models of protracted Andean margin (Fig. 3 e, f) (Slagstad et al., 2020; Slagstad et al., 2017; Slagstad et al., 2013) or was dipping westwards in the model of suturing along the Mylonite Zone at c. 990 Ma (Fig. 3 c) (Brueckner, 2009; Möller and Andersson, 2018; Petersson et al., 2015b). (ii) The Sirdal
1162 magmatic belt represents syn-collision magmatism, therefore recording ongoing continent-
1163 continent collision between 1065 and 1020 Ma (Fig. 3 d).

1164 **4.3 Significance of massif-type anorthosite plutonism**

1165 Massif-type anorthosite plutons formed on Earth only in the Proterozoic. This peculiarity is
1166 inferred to relate directly or indirectly to the secular evolution of the temperature of the
1167 asthenosphere (Ashwal, 1993). The geodynamic context and petrogenesis of AMC plutonism
1168 remain controversial (Ashwal, 1993; Bédard, 2010; Duchesne et al., 1985; Emslie, 1985;
1169 Vander Auwera et al., 2011).

1170 Petrologically, the AMC suite of Rogaland (Fig. 5; Fig. 14) can be accounted for by
1171 differentiation of several parental magmas ranging in composition from high-alumina basalt
1172 (anorthosite plutons) to ferro-basalt (Bjerkreim-Sokndal intrusion and jotunites) in anhydrous
1173 and reduced (QFM to QFM-1) conditions (Charlier et al., 2010; Duchesne and Wilmart, 1997;
1174 Duchesne et al., 1989; Robins et al., 1997; Vander Auwera and Longhi, 1994).

1175 Here, we draw the attention to the fact that the AMC complex is almost entirely devoid of
1176 water-bearing minerals (Duchesne and Charlier, 2005; Longhi et al., 1999). Amphibole
1177 appears only very locally as a late-stage replacement mineral. The dry nature of the magmas
1178 as well as the water-poor to water-absent assemblages of the granulite-facies wall rock of the
1179 AMC plutons (Blereau et al., 2017; Drüppel et al., 2013; Laurent et al., 2018b) are objectively
1180 irreconcilable with the definition of magmatism in a supra-subduction setting (Grove et al.,
1181 2006). Supra-subduction magmatism is induced by fluids released from and fluxing above a
1182 subducting oceanic plate. It typically contains 1–6 wt % H₂O (Plank et al., 2013; Sobolev and
1183 Chaussidon, 1996; Wallace, 2005) and produces hornblende-bearing cumulates (Jagoutz and
1184 Schmidt, 2013). Therefore, in our opinion, models framing AMC magmatism in a supra-
1185 subduction setting (Fig. 3 e, f) (Bybee et al., 2014; Slagstad et al., 2013) are not realistic.

1186 **4.4 Sveconorwegian orogenic plateau**
1187 Evidence for the presence of a past orogenic plateau in Proterozoic orogens is largely indirect (Jamieson and Beaumont, 2013; Rey et al., 2001; Rivers, 2008, 2012; Vanderhaeghe, 2012).
1188
1189 Today, following extension (collapse), exhumation and erosion, the Sveconorwegian orogen exposes widespread gneiss complexes characterized by ductile deformation accompanied by partial melting, compressional structures, and protracted upper amphibolite- to granulite-facies metamorphism, structurally overlain by discontinuous exposures of low-grade supracrustal rocks (Fig. 4). The supracrustal rocks are greenschist- to epidote-amphibolite-facies metavolcanic and metasedimentary sequences, exhibiting partially preserved primary structures and stratigraphic relationships. The age distribution of rocks in the supracrustal complexes matches that in the gneiss complexes. Transition between high-grade and low-grade rock occurs over short distances.
1190
1191
1192
1193
1194
1195
1196
1197
1198

We interpret the gneiss complexes and supracrustal complexes as remnants of the infrastructure and superstructure of an orogenic plateau, respectively, now tectonically juxtaposed along extensional shear zones. Characterization of the geometry, kinematics and geochronology of these shear zones is still very fragmentary today. However, recent data support diffuse late-Sveconorwegian extensional tectonics (Persson-Nilsson and Lundqvist, 2014; Torgersen et al., 2018; Viola et al., 2011).
1199
1200
1201
1202
1203
1204
1205
1206
1207
1208
1209
1210
1211

The sedimentary rocks in the supracrustal complexes offer a window into the surface environment at the time of deposition. As reviewed above, the supracrustal rocks deposited between 1280 and 1050 Ma reflect continental (above sea level) conditions, with evidence for sediment accumulation in fault-bounded intermontane extensional basins (Bingen et al., 2003; Köykkä, 2011; Lamminen, 2011; Spencer et al., 2014). Gneiss complexes in the Bamble and Kongsberg lithotectonic units were exhumed to upper-crustal level after the early-Sveconorwegian orogenic phase (1150–1120 Ma) and, therefore, they can be regarded as part of the orogenic superstructure during the main Sveconorwegian orogeny (after 1065 Ma).
1212
1213
1214
1215
1216
1217
1218
1219
1220
1221
1212 Plutons, produced by partial melting of the lower and middle crust, can be anticipated to
1213 accumulate mainly at the transition between ductile and brittle crust (Brown, 2013). In an
1214 orogenic plateau, they will accumulate between the infrastructure and superstructure. The
1215 Sveconorwegian orogen exposes Sveconorwegian plutons increasing in abundance westwards
1216 and mainly hosted in gneiss complexes (Fig. 5). Plutons intruded between 1065 and 920 Ma,
1217 define a consistent pressure of intrusion of 0.4–0.5 GPa (Table 1). This suggests a rather
1218 constant depth of c. 16 km for the boundary between the infrastructure and superstructure,
1219 through time during the main Sveconorwegian orogeny. Additionally, this is consistent with a
1220 model of stable orogenic plateau extending over large areas in the orogen.

1221 **4.5 End of convergence and collapse of the orogenic plateau**

1222 The switch between plate convergence and plate divergence is a fundamental parameter of
1223 orogeny. However, it is not trivial to constrain in time, because evidence for compression or
1224 extension are distinct in the infrastructure and superstructure of an orogenic plateau. The last
1225 undisputable evidence for convergence in the Sveconorwegian orogen corresponds to eclogite
1226 facies metamorphism dated at 988 ± 6 Ma in the Eastern Segment (Möller et al., 2015).
1227 Several observations, however, indicate that compression continued after this point in the
1228 middle crust (infrastructure), probably to at least c. 930 Ma. (i) In the Eastern Segment, the
1229 internal section and the eclogite-bearing ductile nappe are folded by east-verging to
1230 recumbent folds and later upright folds, recording continued high-grade E–W contraction
1231 (Möller and Andersson, 2018; Möller et al., 2015; Piñán-Llamas et al., 2015; Tual et al.,
1232 2015). Zircon carries a record of these events between c. 978 and 961 Ma. (ii) In the
1233 Telemarkia lithotectonic unit, plutons of the HBG suite exhibit a petrofabric, which is largely
1234 controlled by wall-rock ductile deformation during emplacement (Bolle et al., 2018). A study
1235 of the anomaly of magnetic susceptibility (AMS) of the Holum, Kleivan, and Sjelset plutons
1236 in the Agder area provided evidence for regional E–W compression during intrusion, at 957 ±
7, 936 ± 1 and 932 ± 1 Ma respectively (Fig. 5) (Bolle et al., 2010; Bolle et al., 2003b; Bolle et al., 2018). (iii) In the frontal wedge of the orogen, dykes attributed to the c. 980–945 Ma Blekinge-Dalarna dolerite swarm (Fig. 5) are known to be displaced along discrete ductile shear zones with top-to-east reverse sense of shear. This suggests that thrusting along the Sveconorwegian front took place as late as after c. 945 Ma (Stephens and Wahlgren, 2020a; Wahlgren et al., 1994).

In contrast with this evidence, dykes and sills intruded along brittle structures suggest coeval extension in the upper crust (superstructure). (i) In the Idefjorden lithotectonic unit, WNW–ESE trending mafic to felsic intrusions suggest a phase of NNE-SSW extension between c. 951 and 915 Ma (Fig. 5) (Årebäck et al., 2008; Hellström et al., 2004; Scherstén et al., 2000; Wahlgren et al., 2015). (ii) In the internal section of the Eastern Segment, pegmatite dykes crosscutting the gneiss fabric suggest relaxation between c. 961 and 934 Ma (Andersson et al., 1999; Möller et al., 2007; Möller and Söderlund, 1997; Söderlund et al., 2008b; Söderlund et al., 2002) (iii) In the frontal wedge and the foreland of the orogen, the N-S trending Blekinge-Dalarna dolerites document a phase of E–W extension between c. 978 and 946 Ma (Fig. 5) (Gong et al., 2018; Ripa and Stephens, 2020d; Söderlund et al., 2005).

This cumulatively suggests that the Sveconorwegian orogenic plateau was sustained and grew eastwards until c. 930 Ma, in an overall convergent orogen. Evidence of compression in the ductile middle crust (infrastructure) to c. 930 Ma contrasts with evidence for extension in the same time interval in the brittle upper crust (superstructure), and in the brittle foreland of the orogen.

## 5 Review of Sveconorwegian orogenic models

In light of the evidence summarised and discussed above, we now review and discuss the orogenic models sketched in Fig. 3 are discussed in more detail in the following.
1261 **5.1 Early-Sveconorwegian collision-accretion with suture in Bamble-Kongsberg**
1262 The oldest known Sveconorwegian high-grade metamorphism (1150–1120 Ma) is recorded in
1263 the Kongsberg and Bamble lithotectonic units, in the centre of the Sveconorwegian orogen.
1264 This metamorphism could be interpreted to reflect crustal thickening during an early-
1265 Sveconorwegian collision. This interpretation leads to the conceptual model of Fig. 3 a
1266 involving collision or accretion of an exotic Telemarkia microcontinent to the Idefjorden
1267 lithotectonic unit between 1150 and 1120 Ma, closing an intervening ocean and forming the
1268 Bamble-Kongsberg orogenic wedge (Bingen et al., 2008c; Bingen et al., 2005). At least two
1269 arguments rule out the closure of an oceanic realm. (i) The Mesoproterozoic magmatism
1270 exhibits a significant age overlap between the Bamble–Kongsberg, Telemarkia and Idefjorden
1271 lithotectonic units. Specifically, the 1520–1480 Ma magmatic suites, which are prominent in
1272 the Telemarkia lithotectonic unit, extend well into the Bamble, Kongsberg and Idefjorden
1273 lithotectonic units, thus representing a stitching element of these units around c. 1500 Ma
1274 (Fig. 7). (ii) The granulite-facies low-K calc-alkaline Tromøy Complex in Bamble was
1275 formerly interpreted as an early-Sveconorwegian, c. 1200 Ma old, oceanic volcanic arc
1276 (Andersen et al., 2004a; Andersen et al., 2002b; Knudsen and Andersen, 1999). However,
1277 new data demonstrate that the magmatic protolith of the Tromøy Complex is Gothian (1575 ±
1278 44 to 1544 ± 14 Ma) (Bingen and Viola, 2018), meaning that no evidence for remnants of
1279 early-Sveconorwegian oceanic lithosphere is known in the Bamble lithotectonic unit. There is
1280 therefore no actual geological support for the conceptual model sketched in Fig. 3 a.

1281 **5.2 Early-Sveconorwegian wrench tectonics**
1282 The Bamble and Kongsberg lithotectonic units have been referred to as shear belts in the
1283 literature mostly because of widespread, steep shear foliation zones and penetrative
1284 lithological banding (Starmer, 1991). This intense deformation has inspired tectonic models
1285 (Fig. 3 b) involving long distance early-Sveconorwegian strike-slip transport of the
1286 Telemarkia lithotectonic unit relative to the Idefjorden lithotectonic unit, at the margin of
1287 Fennoscandia, generating a Bamble-Kongsberg transpressional shear belt (Andersen et al.,
1288 2004a; Bingen et al., 2008c; deHaas et al., 1999; Lamminen and Köykkä, 2010). However,
1289 recent field data from the tectonic boundaries between the Bamble, Kongsberg, Telemarkia
1290 and Idefjorden lithotectonic units, and from the centre of the Bamble and Kongsberg
1291 lithotectonic units (Bingen and Viola, 2018; Henderson and Ihlen, 2004; Scheiber et al., 2015)
1292 highlight orthogonal compression and rule out significant wrench tectonics, thus excluding
1293 orogen-scale strike-slip transport. A component of sinistral strike-slip shearing is indeed
1294 recorded by some of the mylonite zones within the Kongsberg lithotectonic unit (Scheiber et
1295 al., 2015). These are, however, compatible with transpressional deformation ensuing only
1296 after the peak of orthogonal deformation and high-grade metamorphism (1150-1120 Ma).

1297 **5.3 Collisional orogeny with suture along the Mylonite Zone**

1298 The Mylonite Zone is a major Sveconorwegian east-southeastward-verging shear zone,
1299 juxtaposing the Eastern Segment beneath the Idefjorden lithotectonic unit. The geological
1300 records of these two units are significantly distinct and, as a consequence, several authors
1301 have argued that the Mylonite Zone may represent a suture zone. An oceanic domain would
1302 have closed at c. 990 Ma between the Eastern Segment, representing the Fennoscandia
1303 continent as lower plate, and distal terranes formed outboard of the Fennoscandia margin in
1304 the west as upper plate (the four western lithotectonic units of the orogen named together
1305 ‘Sveconorwegia’; Fig. 3 c) (Andersson et al., 2002a; Austin Hegardt et al., 2005; Brueckner,
1306 2009; Cornell et al., 2000; Möller and Andersson, 2018; Möller et al., 2015; Petersson et al.,
1307 2015b). This model envisions the pre-990 Ma (pre-collision) magmatism and metamorphism
1308 west of the Mylonite Zone as formed in a supra-subduction setting, above a west-dipping
1309 subduction zone. At least four arguments support this model. (i) The magmatic records in the
1310 Eastern Segment and in the Idefjorden lithotectonic unit are distinct (Fig. 7). Magmatic suites
1311 do not extend across the Mylonite Zone. (ii) Hallandian metamorphism between 1465 and

1312 1385 Ma is documented only east of the Mylonite Zone (Fig. 6) (Söderlund et al., 2002;

1313 Ulmius et al., 2015). (iii) The Sveconorwegian metamorphism in the Eastern Segment reached

1314 eclogite-facies conditions at c. 990 Ma (Möller et al., 2015), significantly after granulite-

1315 facies metamorphism in the Idefjorden hanging wall at c. 1050 Ma (Söderlund et al., 2008a).

1316 Eclogite-facies metamorphism could record continental burial after closure of an ocean basin

1317 (Möller and Andersson, 2018; Möller et al., 2015). (iv) The Lu-Hf isotopic signature of

1318 magmatic rocks in the 1780–1480 Ma interval documents a geochemical disconnect across

1319 the Mylonite Zone, with an average εHf = +3.0 in the Eastern Segment at 1700 Ma against

1320 +4.8 in the Idefjorden lithotectonic unit at 1570 Ma (Petersson et al., 2015a; Petersson et al.,

1321 2015b) (Fig. 8). This difference implies a lower contribution of old continental crust in the

1322 genesis of the magmatic rocks in the Idefjorden lithotectonic unit.

1323 These four pro-arguments, however, are balanced by counterarguments. Specifically (i) the

1324 Orust dolerites (1457 ± 6 Ma) in the Idefjorden lithotectonic unit (Åhäll and Connelly, 1998)

1325 overlap in age with 1465–1385 Ma Hallandian granitic to charnockitic plutonism in the

1326 Eastern Segment. (ii) The Lu-Hf isotopic signature of early-Sveconorwegian magmatism

1327 between 1225 and 1180 Ma is distinctly supra-chondritic in both the Eastern Segment

1328 (bimodal magmatism along the Sveconorwegian front; +1.2 < εHf < + 6.6) and the Telemarkia

1329 lithotectonic unit (Vråvatn Complex; +9 < εHf < + 10; Fig. 8) (Andersen et al., 2007;

1330 Petersson et al., 2015a; Söderlund et al., 2005). This signature attests to coeval depleted

1331 mantle derived magmatism on both side of the Mylonite Zone before the presumed ocean

1332 closure at 990 Ma. (iii) The Mylonite Zone (or geological units in its direct proximity) does

1333 not contain any remnants or slivers of pre- to early-Sveconorwegian (1340–1080 Ma) marine

1334 sedimentary sequences, oceanic lithosphere, oceanic volcanic arc, or ultramafic rocks, such

1335 that no suture zone can be directly constrained.
1336 To conclude, closure of an oceanic basin along the Mylonite Zone at c. 990 Ma represents
1337 a plausible model (Fig. 3 c) (Möller and Andersson, 2018). However, the evidence is not
1338 conclusive at this point of research. In the following text, we do not select this model as the
1339 most probable.

1340 **5.4 Non-collisional (Andean type) orogeny**

1341 In the non-collisional (Andean type) orogenic models (Fig. 3 e, f), the Sveconorwegian
1342 orogen represents an active margin of Fennoscandia, evolving from at least 1280 Ma to after
1343 900 Ma, above an oceanic plate subducting to the east into a trench situated to the west of the
1344 exposed orogen (Falkum and Petersen, 1980; Slagstad et al., 2013). The geological record in
1345 the Sveconorwegian orogen is explained by changes in the conditions of subduction, such as
1346 trench position, subduction angle, convergence rate, convergence direction and age of the
1347 oceanic lithosphere. This model is an adaptation of the tectonic switching model (Collins,
1348 2002; Haschke et al., 2002), which is based on the observation that a retreating or steepening
1349 oceanic subduction is associated with an extensional tectonic regime and abundant
1350 magmatism in the (supra-subduction) upper plate, while an advancing or flattening subduction
1351 is associated with compression, metamorphism and magmatic quiescence.

1352 Different versions of the non-collisional (Andean type) model have been proposed by
1353 Slagstad et al. (2020; 2018; 2017; 2013) and Granseth et al. (2020). These models offer an
1354 elegant and flexible framework for the orogeny. However, we think that they are
1355 irreconcilable with a number of key features and concepts. (i) In its simple expression, the
1356 tectonic switching model predicts either extension or compression in the upper (supra-
1357 subduction) plate. During the main Sveconorwegian orogeny, voluminous magmatism (mafic
1358 and felsic) in the Telemarkia lithotectonic unit would indicate a retreating subduction trench
1359 between 1065 and 1020 Ma, while HP granulite facies metamorphism in the Idefjorden
1360 lithotectonic unit (Söderlund et al., 2008a) would indicate an advancing trench in the same
1361 time interval, in contradiction with the model. (ii) Conceptually, an eastwards oceanic
1362 subduction to the west of the orogen can hardly represent the driving force for westwards
1363 underthrusting of the Eastern Segment at c. 990 Ma to eclogite-facies conditions. Considering
1364 the presumably weak rheology of the lithosphere in the Telemarkia lithotectonic unit around
1365 990 Ma, it is unlikely that compressive stresses from a plate subducting west of the orogen
1366 could be effectively transmitted at least 400 km to the east to the Eastern Segment. (iii) The
1367 Sveconorwegian magmatism (1065–915 Ma) does not represent typical volcanic arc
1368 magmatism. The geochemical signature and petrology of magmatic suites can be related to
1369 lower crustal sources and partial melting conditions (Granseth et al., 2020; Vander Auwera et
1370 al., 2008; Vander Auwera et al., 2011), rather than to an active subduction. The geochemical
1371 signature of the Sirdal magmatic belt (1065–1020 Ma) can be interpreted in both a collisional
1372 setting or a supra-subduction setting (see above). The magmatism between 985 and 915 Ma
1373 lacks a subduction signature and the dry nature of AMC magmatism is not compatible with a
1374 supra-subduction setting (see above). The magmatism between 935 and 915 Ma is exposed
1375 over a zone at least 350 km wide, much larger than a typical volcanic arc. The geographical
1376 polarity of the magmatism in the 935–915 Ma time interval, involving dry plutonism of the
1377 AMC suite in the west and water-bearing plutonism of the HBG suite in the east is opposite to
1378 what should be expected from an east dipping subduction system.

1379 For these different reasons, we remain sceptical that an oceanic subduction in the
1380 hinterland of the orogen could have steered tectonic forces and magmatism inside the orogen
1381 during the main Sveconorwegian orogeny (1065–900 Ma). The non-collisional models
1382 proposed by Slagstad et al. (2020; 2018) omit to propose specific tectonic driving forces,
1383 either oceanic subduction or continental subduction-delamination inside the orogen, to explain
1384 the metamorphism with high-pressure signature in the Bamble–Kongsberg lithotectonic units
1385 (1150–1120 Ma), Idefjorden lithotectonic unit (c. 1050 Ma) and Eastern Segment (c. 990 Ma).
1386 In a recent version of the non-collisional model (Slagstad et al., 2020), the margin of
1387 Fennoscandia is proposed to have been fragmented (into micro-continents) by extension
1388 before c. 1150 Ma and re-amalgamated during the Sveconorwegian orogeny between c. 1150
1389 and 980 Ma to form the Sveconorwegian orogen. However, as discussed previously in this
1390 chapter, there is no evidence between the lithotectonic units for (i) marine sediment sequences
1391 that could represent marine basins, (ii) ophiolites or oceanic volcanic arcs that could represent
1392 oceanic basins, or (iii) ultramafic bodies that could represent exhumed hyperextended
1393 domains. As noted earlier, the low-K calc-alkaline Tromøy Complex in Bamble (Andersen et
1394 al., 2004a) should not be interpreted as an early-Sveconorwegian oceanic volcanic arc
1395 (Bingen and Viola, 2018).

1396 **5.5 Collisional orogeny with suture west of the orogen**

1397 The Grenville orogen is the archetypal example of a large (> 600 km wide) and hot
1398 Mesoproterozoic collisional orogen (Fig. 1) (Gower et al., 2008; Jamieson and Beaumont,
1399 2013; Rivers, 2008, 2012). The Grenville orogeny was long lived (> 110 Myr). It propagated
1400 from a weak Proterozoic lithosphere into the cratonic Archean foreland, with thrusting along
1401 two orogen parallel, continuous, crustal-scale shear zones (the Allochthon Boundary Thrust
1402 and Grenville Front) and two main phases of orogenic convergence (the Ottawan, 1090–1020
1403 Ma, and Rigolet, 1010–980 Ma, phases; Fig. 1). Protracted high-temperature–low-pressure
1404 high-grade metamorphism in the hinterland, associated with both crustal- and mantle-derived
1405 magmatism, was coeval with comparatively short-lived high-pressure metamorphism in thick
1406 thrust slices towards the foreland (Groulier et al., 2018a; Indares, 2020; Rivers, 2008). The
1407 first order architecture of the Sveconorwegian orogen and its orogenic evolution are
1408 comparable to that of the Grenville orogen (Cawood and Pisarevsky, 2017; Gower, 1985;
1409 Gower et al., 2008; Hoffman, 1991; Rivers, 2008, 2012). An analogy in the geodynamic
1410 evolution is therefore natural.
1411 In the collisional models (Fig. 3 d), closure of (one or several) oceanic basin(s) to the west
1412 of the exposed orogen was followed by the collision of Baltica (Fennoscandia) with (one or
1413 several) continental plate(s) at and after 1065 Ma (Bingen et al., 2008c; Bogdanova et al.,
1414 2008; Cawood and Pisarevsky, 2017; Gower et al., 2008; Ibanez-Mejia et al., 2011; Li et al.,
1415 2008; Pisarevsky et al., 2014; Stephens and Wahlgren, 2020b; Weber et al., 2010).
1416 Consumption of these oceanic basins involved subduction and formation of volcanic arcs,
1417 either at the margin of Fennoscandia or in an outboard position prior to final collision. During
1418 collision, the Sveconorwegian orogen was situated in the upper plate position, on the
1419 Fennoscandia side of the main suture zone.

1420 Here are some key features supporting the collisional model for the main Sveconorwegian
1421 orogeny (1065–900 Ma). (i) The Sveconorwegian orogen is c. 550 km wide, i.e. wider than
1422 any present-day Andean orogen. It exhibits a c. 550 km wide zone of convergent tectonics
1423 (1065–930 Ma) and a c. 350 km wide zone of syn-orogenic magmatism (Fig. 5). (ii) The
1424 orogen has the structure of an extended (collapsed) orogenic plateau, with the juxtaposition of
1425 high-grade gneiss complexes representing a middle crustal infrastructure, against low-grade
1426 supracrustal rocks representing a brittle superstructure, and plutons representing the product
1427 of lower- to middle-crustal melting during orogeny (Fig. 4) (Andersen et al., 2001; Granseth
1428 et al., 2020; Vander Auwera et al., 2011). The gneiss complexes carry evidence for protracted
1429 (> 110 Myr) middle-crustal high-temperature-low-pressure metamorphism (Bingen et al.,
1430 2008b; Blereau et al., 2017; Laurent et al., 2018a; Laurent et al., 2018b; Slagstad et al., 2018).
1431 (iii) High pressure granulite- and eclogite-facies rocks attest to crustal thickening, up to c. 70
1432 km, between c. 1050 and 990 Ma (Möller and Andersson, 2018; Söderlund et al., 2008a).
1433 After peak metamorphism, these rocks were incorporated and overprinted into the middle-
1434 crustal infrastructure. They were probably more abundant in the orogen than what is apparent
1435 from their exposure. (iv) The orogenic zone grew towards the foreland, in a stepwise fashion
1436 with time. This process involved thrusting along crustal scale shear zones (Mylonite Zone and Sveconorwegian front). This pattern is typical of collision orogens (Royden et al., 2008). (v)
1437
1438 The orogenic zone lacks evidence for syn-orogenic marine sedimentary sequences, in spite of a largely exposed superstructure (Fig. 4), and therefore was above sea-level during the entire orogeny.
1439
1440
1441

<figure><img src="image_30.png" /></figure>

1442
1443 **Figure 16.** Common Pb isotopic composition in the $^{206}\text{Pb}/^{204}\text{Pb}$ vs. $^{207}\text{Pb}/^{204}\text{Pb}$ diagram, with reference growth curves of terrestrial common Pb ($\mu_{(2)}$ = 9.74 (Stacey and Kramers, 1975) and
1445 $\mu_{(2)}$ = 9.5, with $\mu = {}^{238}\text{U}/{}^{204}\text{Pb}$). (a) Compilation of data from the Sveconorwegian orogen.

1446 Highlighted symbols represent initial ratio of plutonic suites dated between 1150 and 930 Ma.

1447 Initial ratio (ratio corrected for U decay since intrusion) is calculated for analyses of K-

1448 feldspar, plagioclase, orthopyroxene and whole-rock with $\mu < 3$. The initial ratio of plutonic

1449 rocks defines a short trend below the reference growth curve of Stacey and Kramers (1975).

1450 High alumina orthopyroxene megacrysts (HAOM) hosted in the anorthosite plutons are

1451 situated at the radiogenic (upper-right) end of the trend. They are interpreted to represent a

1452 mafic, mantle-derived, underplate, formed at c. 1040 Ma and remelted at c. 930 Ma. The

1453 granite plutons partly sourced from metasedimentary protoliths, like the Flå and Bohus

1454 muscovite-bearing plutons, are situated at the less radiogenic (lower-left) end of the trend.

1455 The hornblende-biotite granite plutons, ranging from c. 1150 to 930 Ma, sourced from

1456 metaigneous protoliths, cluster in the centre of the trend (Andersen, 1997; Andersen et al.,

1457 2001; Andersen et al., 1994; Andersen and Munz, 1995; Bingen et al., 1993; Vander Auwera

1458 et al., 2014a; Weis, 1986). (b) Compilation of initial isotopic compositions for Baltica

1459 (Fennoscandia) and selected late-Mesoproterozoic orogenic belts. Data for the

1460 Sveconorwegian orogen are copied from panel (a). A variety of ortho- and paragneisses from

1461 the hinterland of the Grenville orogen, including the Composite Arc (Ontario), Frontenac-

1462 Adirondack (Ontario) and Llano uplift (Texas) overlap with the data of the Sveconorwegian

1463 orogen, as well as orthogneisses from the Oaxaquia lithotectonic unit (Mexico) (Cameron et

1464 al., 2004; DeWolf and Mezger, 1994). Granite plutons (950 Ma) in the Aguapei Belt in

1465 Amazonia are characterized by a more radiogenic $^{207}\text{Pb}$ signature than coeval rocks in the

1466 Sveconorwegian orogen, consistent with involvement of an older Paleoproterozoic basement

1467 in this orogen (Geraldes et al., 2001). Data for rapakivi granite plutons from Fennoscandia on

1468 Paleoproterozoic and Archean basement, as well as Paleoproterozoic galena deposits from the
1469 Svecokarelian orogen are shown for reference (Andersson et al., 2002b; Rämö, 1991;
1470 Vaasjoki, 1981).

1471 **5.6 Conjugate margins in Rodinia**

1472 In classical Rodinia assembly models, Laurentia and Baltica were probably already
1473 contiguous at low latitudes at c. 1260 Ma as part of Nuna (Columbia), facing an ocean, the
1474 Mirovoio ocean (Buchan et al., 2000; Evans and Mitchell, 2011; Pisarevsky et al., 2014;
1475 Zhang et al., 2012). Opening of the Asgard sea (north of Baltica; Fig. 1), clockwise rotation
1476 and drift of Baltica relative to Laurentia, and consumption of the Mirovoio ocean (south of
1477 Baltica) led to collision of Amazonia with Laurentia and Baltica, involving three sequential
tectonic phases (Bogdanova et al., 2008; Cawood and Pisarevsky, 2017; Gower et al., 2008;
1478 Hynes and Rivers, 2010; Ibanez-Mejia et al., 2011; Johansson, 2009; Li et al., 2008;
1479 Pisarevsky et al., 2014; Roberts, 2013; Tohver et al., 2004a; Weber et al., 2010). (i) Collision
1480 between Amazonia and the southwestern part of Laurentia starting at c. 1200 Ma and
generating the Llano section of the Grenville orogen and the Sunsás orogen (Fig. 1). (ii)
1481 Sinistral transpression between Amazonia and Laurentia, between c. 1150 and 1050 Ma,
1482 generating the Grenville orogen. (iii) Collision between Amazonia and Baltica at c. 1060 Ma,
1483 following closure of the intervening oceans, producing the Sveconorwegian and Putumayo
1484 orogens (Fig. 1) (Boger et al., 2005; Ibanez-Mejia et al., 2011; Tohver et al., 2004b; Tohver et
1485 al., 2005).
1486
1487

1488 Following Cawood et al. (2010), a subduction system was initiated along the northern open
1489 margin of Rodinia (Asgard sea; Fig. 1) after the Amazonia–Laurentia–Baltica collision (i.e.
after 1000 Ma). In this model, Tonian sediments sequences and volcanic and plutonic rocks
1490 hosted in variably far-travelled nappes of the Caledonides of NE Greenland, Scandinavia,
1491 Svalbard and Scotland (Augland et al., 2014; Cawood and Pisarevsky, 2017; Cawood et al.,
1492 2015; Corfu, 2019; Cutts et al., 2009; Kalsbeek et al., 2000; Kirkland et al., 2006, 2007) are
1494 interpreted as fragments of an accretionary orogen, the Valhalla orogen, at the margin of
1495 Rodinia (Fig. 1). An orogenic phase, including magmatism, metamorphism, and deformation
1496 (Renlandian) took place between 980 and 910 Ma (Cawood et al., 2010), therefore
1497 overlapping with metamorphism in the Sveconorwegian orogen (Table 1).

1498 Several Mesoproterozoic basement inliers in the Andes of Colombia (Garzón, Las Minas)
1499 and in Mexico (Oaxaquia lithotectonic unit) are characterized by high-grade metamorphism,
1500 dated consistently between 1000 and 980 Ma (Zapotecan-Putumayo orogenies). These
1501 lithotectonic units are interpreted as oceanic volcanic arcs formed in ocean tracts between
1502 Laurentia, Amazonia and Baltica (after c. 1460 Ma) (Fig. 12 a, b) and involved in the
1503 collision zone between these plates (Fig. 1) (Cardona et al., 2010; Cordani et al., 2005;
1504 Ibanez-Mejia et al., 2015; Ibanez-Mejia et al., 2011; Jiménez-Mejía et al., 2006; Keppie et al.,
1505 2003; Keppie and Ortega-Gutiérrez, 2010; Weber and Köhler, 1999; Weber et al., 2010).

1506 In archetypal Rodinia reconstructions (Fig. 1), the hinterland of the Sveconorwegian
1507 orogen is facing Mesoproterozoic basement inliers in the Andes of Colombia (Garzón, Las
1508 Minas) and in Mexico (Oaxaquia lithotectonic unit) and the hinterland of the Grenville
1509 orogen. The isotopic signature of these units is compared in a εNd vs. time diagram (Fig. 13
1510 b). The Quebecia and Telemarkia lithotectonic units, located in the hinterland of the exposed
1511 Grenville and Sveconorwegian orogens respectively, represent coeval continental growth
1512 zones generated by volcanic arc and back arc magmatism between c. 1520 and 1480 Ma
1513 (Pinwarian and Telemarkian phases; Table 1; Fig. 13 b) (Dickin and Higgins, 1992; Groulier
1514 et al., 2018b). These units are characterized by very similar isotopic evolution trends starting
1515 from close to the Depleted Mantle reservoir at and after 1520 Ma and decreasing along a
1516 continental recycling trend to near-chondritic value at c. 1000 Ma. The basement inliers in the
1517 Andes of Colombia (Garzón, Las Minas) and in Mexico (Oaxaquia lithotectonic unit) define
1518 evolution trends starting at c. 1380 Ma and 1300 Ma, respectively, that overlap with the
1519 Telemarkia trend (Ibanez-Mejia et al., 2015; Lawlor et al., 1999; Weber and Köhler, 1999). In
1520 the hinterland of the Grenville Belt, the Composite Arc and Frontenac-Adirondack
1521 lithotectonic units define the most juvenile trend, starting at c. 1380 Ma (Fig. 13 b) (Daly and
1522 McLelland, 1991; Dickin et al., 2010; Marcantonio et al., 1990). These units are interpreted as
1523 marginal or outboard volcanic arcs, back-arcs and microcontinents assembled (or
1524 reassembled) to Laurentia early during the Grenvillian orogeny (Shawinigan phase, 1190-
1525 1140 Ma) (Carr et al., 2000; Hanmer et al., 2000; Rivers, 2008). Interestingly, the 1280-1200
1526 Ma magmatism in the Composite Arc (Elzevirian) has a Nd isotopic signature approaching
1527 that of the Depleted Mantle reservoir (Carr et al., 2000; Corfu and Easton, 1995; Corriveau
1528 and van Breemen, 2000; Dickin and McNutt, 2007), very similar to the one of cœval 1280-
1529 1200 Ma continental magmatism in the Sveconorwegian orogen (Sæsvatn-Valldal bimodal
1530 volcanism) (Brewer et al., 2004).

1531 In the 206Pb/204Pb vs. 207Pb/204Pb diagram (Fig. 16), the initial isotopic composition of
1532 Sveconorwegian plutonic rocks intruded between 1150 and 930 Ma defines a short trend
1533 below the evolution curve of terrestrial common Pb of Stacey and Kramers (1975) (μ(2) =
1534 9.74). Ortho- and paragneisses from the hinterland of the Grenville Belt, including the
1535 Composite Arc (Ontario), Frontenac-Adirondack (Ontario) and Llano uplift (Texas) and
1536 orthogneisses from the Oaxaquia lithotectonic unit (Mexico) overlap with the data of the
1537 Sveconorwegian orogen (Cameron et al., 2004; DeWolf and Mezger, 1994).

1538 To summarize, Nd and Pb isotopic data (Fig. 13; Fig. 16) and detrital zircon data (Fig. 12)
1539 underscore the existence of juvenile lithotectonic units generated at and after c. 1520 Ma
1540 exposed in the hinterland of the Grenville and Sveconorwegian orogens and the basement
1541 inliers in the Andes of Colombia (Garzón, Las Minas) and in Mexico (Oaxaquia lithotectonic
1542 unit). These isotopic data therefore support to join these lithotectonic units in the core of the
1543 collision zone between Laurentia, Amazonia and Baltica, in a classical Rodinia reconstruction
1544 (Fig. 1). These data also suggest that, in the collision model of Fig. 3 d, the continental margin
1545 colliding with the Sveconorwegian orogen possessed a weak lithosphere similar to the one of
1546 the Telemarkia lithotectonic unit (as opposed to a stronger cratonic lithosphere).

1547 **6 Model of large, hot and long-duration continental collision**

1548 The previous discussion argues for a collisional model for the main Sveconorwegian orogeny.
1549 Here, we further develop a model of large, hot and long-duration continent-continent collision
1550 starting at c. 1065 Ma, wherein the five lithotectonic units of the orogen are endemic to
1551 Fennoscandia (Fig. 3 d). The plate tectonic interpretation of the pre-collision evolution
1552 between 1280 and 1080 Ma is still largely speculative. The model is fitted into a classical
1553 Rodinia assembly framework (Fig. 1), involving an Amazonia-Laurentia-Baltica collision, as
1554 discussed previously. However, the model is based upon evidence from within the exposed
1555 Sveconorwegian orogen (and its foreland) and therefore independent of Rodinia models.

1556 -----------------
<figure><img src="image_31.png" /><figcaption>**Figure 17.** Schematic geodynamic model for the 1280–1150 Ma time interval. (a) Sketch map of the Sveconorwegian orogen, with position and list of plutonic and supracrustal complexes. (b, c) Interpretative E-W cross sections of the Sveconorwegian orogen and speculative linkage westwards. The exposed part of the orogen is limited by an arrowed bracket. The limit between lithotectonic units is schematically represented by a white vertical dashed line.

References and explanations in the text.</figcaption></figure>

1557

1558

1559

1560

1561

1562

1563

1564

------------------
<figure><img src="image_32.png" /><figcaption>**Figure 18.** Schematic geodynamic model for the 1150–1080 Ma time interval. (a) Sketch map of the Sveconorwegian orogen, with distribution of metamorphism, magmatism and clastic sediment basins. (b) Pressure-temperature diagram with fields of the main metamorphic facies following Spear (1993): E: eclogite facies, G: granulite facies, A: amphibolite facies. (c, d) Interpretative E-W cross sections of the Sveconorwegian orogen and speculative linkage westwards. Explanations and references in the text.</figcaption></figure>

1565

1566

1567

1568

1569

1570

1571

1572

-----------------
1573

## 6.1 1280-1080 Ma, pre-collision: lithospheric mantle delamination

1574

The geological record for the pre- to early-Sveconorwegian 1280-1080 Ma time interval is
very distinct in the Telemarkia, Kongsberg-Bamble and Idefjorden lithotectonic units.

1575

1576

Abundant bimodal magmatism between 1280 and 1145 Ma, and protracted (upper) crustal extension in Telemarkia between 1280 and 1080 Ma (Fig. 17; Fig. 18) contrast with

1577

1578

amphibolite- to granulite-facies metamorphism and shortening in the Kongsberg-Bamble

1579

lithotectonic units between 1150 and 1120 Ma (Fig. 18). Except for a few dolerite dykes, the

1580

Idefjorden lithotectonic unit is lacking evidence for magmatism, metamorphism and

1581

deformation between 1280 and 1080 Ma, and therefore it is regarded as having played the role
of a passive buttress during this time interval. There is no evidence for closure of marine or
oceanic basins between the Telemarkia, Kongsberg-Bamble and Idefjorden lithotectonic units

1582

1583

1584

(Scheiber et al., 2015). These different features cannot be explained by a simple model of

1585

regional scale inversion from extension to compression at c. 1150 Ma, as one would anticipate
compression structures to be located in the weakest Telemarkia lithosphere, or distributed

1586

evenly throughout the Telemarkia, Bamble and Kongsberg lithotectonic units.

1587

1588

In Fig. 17, we propose that upwelling of asthenosphere and development of an orogenic
plateau started at c. 1280 Ma in the Telemarkia lithotectonic unit. Repeated pulses of bimodal
magmatism between 1280 and 1145 Ma provide evidence for upwelling and decompression
melting of asthenospheric mantle (Fig. 7). The most prominent mafic volcanic rocks (Fig. 11;
Sæsvatn-Valldal, Nissedal, Morgedal and Gjuve metabasalts) exhibit a within-plate
geochemical signature and supra-chondritic Nd isotopic signature (+2.6 < εNd < +6.3)

1589

1590

1591

1592

1593

1594

implying sourcing in the asthenosphere (Fig. 13 a) (Brewer et al., 2002; Brewer et al., 2004;
Spencer et al., 2014). The voluminous felsic gneisses of the Vråvatn Complex (1220-1190

1595

1596

Ma) have Hf isotopic signature of zircon (+9 < εHf < +10) also close to the depleted mantle

1597

reservoir at 1210 Ma (εHf = +12) (Fig. 8) (Andersen et al., 2007). These values indicate that
1598 the Vråvatn Complex was not produced principally by partial melting of the Telemarkian
1599 (1520–1480 Ma) crust. Rather, it was probably produced by partial melting of a mafic lower
1600 crust or mafic underplate, itself produced shortly before in the depleted mantle (a maximum
1601 of some 50 Mys before 1210 Ma; Andersen et al., 2007). The earliest magmatism between
1602 1280 and 1190 Ma occurred in the centre of the Telemarkia lithotectonic unit (Vråvatn,
1603 Nissedal, Sæsvatn–Valldal, Iveland-Gautestdad; Fig. 6; Fig. 17) while younger magmatism
1604 between 1170 and 1140 Ma is more abundant towards the periphery of the lithotectonic unit
1605 (mainly eastwards and southwards) and is well recorded into the Kongsberg and Bamble
1606 lithotectonic units. This geographic distribution suggests that mantle upwelling affected
1607 progressively a larger area between c. 1280 and 1145 Ma (Fig. 17).

1608 Upwelling of hot asthenosphere at c. 1280 Ma induced partial melting at the base of the
1609 crust. This weakening of the lower crust possibly initiated decoupling between the crust and
1610 the lithospheric mantle and progressive delamination of the lithospheric mantle between c.
1611 1280 and 1145 Ma. Alternatively, protracted upwelling of asthenosphere between 1280 and
1612 1145 Ma progressively induced convective removal (or displacement) of the continental
1613 lithospheric mantle. Both interpretations resulted in uplift, formation of a plateau and
1614 extension in the crust (Dewey, 1988; Li et al., 2016).

1615 Evidence for an orogenic plateau involving uplift and extension in the upper crust is
1616 provided by the sedimentology of low-grade sedimentary rocks deposited between 1260 and
1617 1080 Ma in Telemarkia. The sediments of the Bandak succession are high-energy immature
1618 deposits, accumulated in continental (above sea level) intermontane basins (Bingen et al.,
1619 2003; Köykkä, 2011; Lamminen, 2011; Spencer et al., 2014). The limited lateral extent of the
1620 basin infills, the existence of at least two major internal unconformities in the Bandak
1621 succession, and the direct evidence for normal syn-sedimentary growth faults, suggest active
1622 extension during accumulation (Fig. 11; Fig. 12) (Laajoki, 2002; Laajoki et al., 2002;
1623 Lamminen, 2011). This orogenic plateau does not satisfy to the definition of a Tibetan orogenic plateau, as evidence for crustal thickening and protracted metamorphism is lacking in the 1280-1080 Ma time interval inside the Telemarkia lithotectonic unit.

1626 Compression in the Kongsberg-Bamble lithotectonic units in the interval between 1150 and 1120 Ma can be explained by compression at the margin of the plateau and foundering (subduction) below the Kongsberg-Bamble lithotectonic units of the lithospheric mantle slab delaminated below Telemarkia (Fig. 18). We suggest that the pull effect of foundering generated subsidence in the crust, high-grade metamorphism (up to 1.15 GPa) and deformation with lithological banding and commonly steep lineation in Kongsberg-Bamble between 1150 and 1120 Ma.

1633 Eventual breakoff of the mantle slab triggered exhumation of the Bamble and Kongsberg lithotectonic units after 1120 Ma (Fig. 18 d). The volumetrically minor alkaline Morkheia monzonite suite, located just north of the Bamble-Telemarkia boundary zone, may record this event with local melting of a sliver of lithospheric mantle at c. 1134-1130 Ma (Fig. 18 c)
1634 (Heaman and Smalley, 1994). Exhumation to upper crustal levels (1105-1080 Ma) was associated with northwestwards thrusting of Bamble and westwards thrusting of the Kongsberg onto Telemarkia and reworking of plutons emplaced shortly before the foundering process (Henderson and Ihlen, 2004; Scheiber et al., 2015). Exhumation of the Bamble and Kongsberg lithotectonic units after 1105 Ma was associated with fluid-rock interaction, albitization and scapolitization (Engvik et al., 2017). Erosion provided the clastic material stored in the Heddal Group and Eidsborg Formation in Telemarkia (Fig. 6; Fig. 11; Fig. 12; Fig. 18).

1645 The plate tectonic context and paleogeographic setting of this pre- to early-
1646 Sveconorwegian asthenosphere upwelling, plateau development and sub-continental
1647 lithospheric mantle delamination model is difficult to assess. The asthenosphere upwelling
1648 could be related to a deep mantle plume, similar to the ones that generated the four mafic

1649 dyke swarms of the Central Scandinavian dolerites in the cratonic center of Fennoscandia

1650 between c. 1271 and 1246 Ma (Brander et al., 2011; Söderlund et al., 2006), and dolerites

1651 along the Sveconorwegian front (Protogine zone dolerites) between 1221 and 1215 Ma

1652 (Söderlund et al., 2005) (Fig. 17). Alternatively, in recent paleogeographic models (Cawood

1653 and Pisarevsky, 2017), the Sveconorwegian orogen (Telemarkia–Bamble–Kongsberg

1654 lithotectonic units) is located in a continent back-arc position on the Fennoscandia side of an

1655 active volcanic arc, in the 1280–1080 Ma interval, during consumption of the oceans between

1656 Baltica, Amazonia and Laurentia (Fig. 17; Fig. 18) (Bingen et al., 2003; Brewer et al., 2002;

1657 Roberts and Slagstad, 2015; Slagstad et al., 2017; Spencer et al., 2014). This arc would be

1658 located to the west of the exposed orogenic belt and possibly disappeared by tectonic erosion

1659 (Spencer et al., 2014). A Cenozoic analogue to this Mesoproterozoic evolution would be

1660 plateau building and lithospheric delamination in the Colorado Plateau and the North

1661 American Cordillera (Bao et al., 2014; Levander et al., 2011).

1662 -----------------
<figure><img src="image_33.png" /><figcaption>**Figure 19.** Schematic geodynamic model for the 1065–1030 Ma time interval. (a) Sketch map of the Sveconorwegian orogen, with distribution of metamorphism, magmatism and clastic sediment basins. (b) Pressure-temperature diagram. (c) Interpretative E-W cross section.

Explanations and references in the text.</figcaption></figure>

1663

1664

1665

1666

1667

1668

---
<figure><img src="image_34.png" /></figure>

1669

1670 **Figure 20.** Schematic geodynamic model for the 1030–1000 Ma interval. (a) Sketch map of

1671 the Sveconorwegian orogen, with distribution of metamorphism and magmatism. (b)

1672 Pressure-temperature diagram. (c) Interpretative E-W cross section. Explanations and
1673 references in the text.

1674

---

1675 **6.2 1065-1000 Ma: main Sveconorwegian continental collision**

1676 After a period of quiescence, the orogenic zone grew substantially around 1065 Ma, both

1677 eastwards (continentwards) and westwards, to include the entire Idefjorden and Telemarkia

1678 lithotectonic units (Fig. 19; Fig. 20). Widespread compressional deformation, high-grade

1679 metamorphism, partial melting and magmatism are recorded in these units between 1065 and

1680 1000 Ma (Agder phase). Little is recorded in the Kongsberg–Bamble lithotectonic units,

1681 which were exhumed to high crustal levels (superstructure) and juxtaposed as reflected by
1682 their current position before 1080 Ma. High-pressure granulite facies metamorphism in the
1683 Idefjorden lithotectonic unit, dated to c. 1050 Ma, contrasts with the voluminous granite
1684 magmatism of the Sirdal magmatic belt (1065-1020 Ma) and low-pressure granulite-facies
1685 metamorphism (1045-990 Ma) culminating at UHT conditions (1030-1005 Ma) in the west
1686 of the Telemarkia lithotectonic unit (Fig. 5; Fig. 19; Fig. 20). The width of the orogenic zone
1687 (minimum of 460 km), the large volume of magmatism with syn- to late-collision
1688 geochemical signature (Sirdal magmatic belt), the paired belts of high-pressure (towards the
1689 foreland) vs. high-temperature (towards the hinterland) metamorphism, and the structural
1690 evidence for convergence suggests that the Sveconorwegian orogeny entered the main phase
1691 of continent-continent collision around 1065 Ma.

1692 In Fig. 19, we propose that collision resulted in the formation of a Tibetan-style orogenic
1693 plateau (Jamieson and Beaumont, 2013) extending from the Telemarkia to the Idefjorden
1694 lithotectonic units. The infrastructure of this orogenic plateau is defined by widespread gneiss
1695 complexes characterized by partial melting, compressive ductile deformation and
1696 amphibolite- to granulite-facies metamorphism between 1050 and 1000 Ma. Evidence from
1697 the superstructure of this plateau is scanty, simply because little upper crustal rocks younger
1698 than 1050 Ma are preserved. The N-S trending Kalhovd Formation consists of unconformable
1699 conglomerate and immature sandstone, deposited after c. 1054 Ma, in a continental (above sea
1700 level) intermontane basin (Fig. 5; Fig. 11; Fig. 19; Fig. 20). This basin was downfaulted along
1701 the Mandal-Ustaoset fault zone, possibly during deposition, recording extension after c. 1054
1702 Ma in the upper crust in the centre of the Telemarkia lithotectonic unit.

1703 Dynamic of the mantle in the collision zone would be simulated by a "pro-plate" (upper-
1704 plate) Tibetan-style delamination numerical models by Li et al. (2016). In Fig. 19, we propose
1705 that mantle upwelling under the Telemarkia lithotectonic unit was counterbalanced by mantle
1706 downwelling and lithospheric mantle delamination and foundering under the Idefjorden
1707 lithotectonic unit. In the Idefjorden lithotectonic unit, the crust was pulled down by a
1708 lithospheric mantle slab to reach peak high-pressure low-temperature granulite facies
1709 conditions (c. 1.2–1.5 GPa, 740–780 °C) between c. 1052 and 1046 Ma (Fig. 19) (Bingen et
1710 al., 2008b; Söderlund et al., 2008a). The crust in the Idefjorden lithotectonic unit was not
1711 affected by orogenic processes before 1050 Ma and therefore could reach high-pressure
1712 conditions before melting. Decoupling between the mantle slab and the crust (breakoff)
1713 probably took place when partial melting reactions were activated in the lower crust. During
1714 exhumation, widespread migmatitization (including muscovite-, biotite- and amphibole-
1715 dehydration melting) is observed at regional scale between c. 1040 and 1000 Ma in the
1716 Idefjorden lithotectonic unit (Fig. 10; Fig. 20; Table 2). Migmatitization took place in
1717 convergent setting; east of the Oslo rift, it is associated with a well-defined top-to-west
1718 direction of transport (Viola et al., 2011). Lamprophyre dykes, close to the boundary between
1719 the Kongsberg and Idefjorden lithotectonic units, attests to local melting of lithospheric
1720 mantle material at c. 1030 Ma (Bingen and Viola, 2018), and is consistent with a model of
1721 foundering of the lithospheric mantle around 1030 Ma.

1722 In the Telemarkia lithotectonic unit (Agder area), the NNW-SSE trending Sirdal magmatic
1723 belt attests to voluminous crustal melting between c. 1065 and 1020 Ma (Fig. 5; Fig. 19)
1724 (Bingen et al., 2015; Coint et al., 2015; Granseth et al., 2020; Slagstad et al., 2013). As
1725 discussed earlier, it contains high-K calc-alkaline quartz-monzonite-granodiorite plutons
1726 associated with minor ultrapotassic rocks (Fig. 7; Fig. 15) (Bingen et al., 1993; Bingen and
1727 van Breemen, 1998a). Such calc-alkaline granitoids can be derived by partial melting of lower
1728 crustal mafic metaigneous rocks. Enrichment in K and other large ion lithophile elements
1729 (LILE) implies either that this crustal source was previously enriched in LILE or that the
1730 melts were mixed with ultrapotassic lamprophyric melts, themselves generated from
1731 lithospheric mantle previously enriched in LILE. The most straightforward interpretation is
that this lithospheric mantle source was part of a mantle wedge enriched in LILE by supra-
subduction fluids between 1520 and 1480 Ma. It would then become part of a subcontinental
lithospheric mantle after 1480 Ma, and finally, it would melt during collision between 1065
and 1020 Ma, heated during orogeny. Two observations support this three stage model: i) the
1065-1020 Ma Sirdal magmatic belt overlaps geographically with the 1520-1480 Ma Suldal
magmatic arc (Roberts et al., 2013); ii) the near-chondritic Nd isotopic signature of the Feda
suite granitoids (-1 < εNd < +1.5) and ultrapotassic enclaves (+1 < εNd < +1.5) are lying on the
evolution vector of the crust generated at 1520-1480 Ma (Fig. 13 a). A similar interpretation is
provided for near-coeval (1063 ± 3 Ma) high-Sr-Ba quartz-monzonite plutons in the Quebecia
lithotectonic unit of the Grenville orogen (Michaud pluton hosted in the c. 1500 Ma Escumins
supracrustal rocks) (Groulier et al., 2018a).

Voluminous melting of the crust clearly requires an appropriate heat source. The high-
alumina orthopyroxene megacrysts (HAOM) and plagioclase megacrysts hosted in the c. 930
Ma anorthosite plutons constrain mafic magmatism at the base of the crust at 1041 ± 17 Ma
coeval with formation of the Sirdal magmatic belt (Bybee et al., 2014; Slagstad et al., 2018;
Vander Auwera et al., 2014b). The supra-chondritic Nd isotopic values for the megacrysts
(+2.8 < εNd(1041 Ma) < +5.3) trace the source of this magmatism to the asthenosphere (Fig. 13
a). Here we suggest that upwelling of hot asthenosphere in the collision zone generated
asthenospheric melts, produced underplates (with HAOM), reheated and destabilized the
lithospheric mantle, and generated minor lamprophyre melts (from this lithospheric mantle)
(Fig. 19; Fig. 20). Heating of the crust produced granitoids of the Sirdal magmatic belt. After
extraction of these melts, protracted heating in the crust resulted in a first phase of granulite-
facies metamorphism (M1), reaching ultra-high temperature conditions (0.7-0.5 GPa, 900-
950 °C) between 1030 and 1005 Ma (Fig. 20) (Blereau et al., 2017; Drüppel et al., 2013;
Laurent et al., 2018b).
1757 The HAOM hosted in the anorthosites record a pressure of crystallization of 1.1 GPa and
1758 an age of 1041 ± 17 Ma (Bybee et al., 2014; Charlier et al., 2010). These numbers imply a
1759 crustal thickness of at least 42 km around 1040 Ma, corresponding to a moderate crustal
1760 overthickening relative to the standard 30 km. The regional folding observed at various scales
1761 in the Telemarkia lithotectonic unit, also requires at least one phase of compression between
1762 1030 and 1000 Ma.

1763 In Fig. 17 to Fig. 20, the Oaxaquia and inliers of in the Andes of Colombia, which contain
1764 evidence of high-grade metamorphism between 1050 and 980 Ma, are represented
1765 speculatively as volcanic arcs in the ocean between, Amazonia Laurentia and Baltica (Fig. 8,
1766 Fig. 12; Fig. 13; Fig. 16). They were first accreted to Amazonia (Putumayo) before colliding
1767 with Baltica around 1050 Ma (Fig. 19; Fig. 20) (Ibanez-Mejia et al., 2011; Lawlor et al.,
1768 1999; Weber et al., 2010). In this framework, the Telemarkia–Kongsberg–Bamble units are
1769 indeed situated in a back-arc domain, in the time interval between 1280 and 1080 Ma, before
1770 closure of all oceans, and they ended-up in upper plate position at c. 1065 Ma during collision
1771 with Amazonia (Fig. 18; Fig. 19; Fig. 20). Upwelling of hot asthenosphere in the collision
1772 zone was possibly promoted by break-off all oceanic lithospheric plates to the west of the
1773 orogen when subductions ceased (Fig. 19).

1774 -----------------
<figure><img src="image_35.png" /></figure>

**Figure 21.** Schematic geodynamic model for the 1000–980 Ma time interval. (a) Sketch map of the Sveconorwegian orogen, with distribution of metamorphism and magmatism. (b) Pressure-temperature diagram. (c) Interpretative E-W cross section. Explanations and references in the text.

## 6.3 1000–920 Ma: long-duration growth of the collision zone

At c. 1000 Ma, orogeny propagated to the east, all the way into the Eastern Segment.

Continued high-temperature low-pressure metamorphism in the west of the collision zone and voluminous magmatism contrast with high pressure metamorphism in the east, suggesting that
1785 the western part was characterized by protracted mantle upwelling while the eastern part by
1786 underthrusting and mantle downwelling. In Fig. 21, Fig. 22 and Fig. 23, we propose that the
1787 Tibetan-style orogenic plateau grew towards the foreland to includes the Eastern Segment and
1788 covered the entire orogen. The Eastern Segment offers a nice example of how crust
1789 underthrust to high-pressure conditions is incorporated into a melt-lubricated middle-crustal
1790 infrastructure of an orogenic plateau. We suggest that the orogenic plateau was sustained to c.
1791 930 Ma before it collapsed.

1792 6.3.1 Eastward growth of orogenic plateau

1793 The Eastern Segment was a cold lithospheric segment of Fennoscandia foreland affinity,
1794 unaffected by Sveconorwegian orogenic processes before 1000 Ma. At this point in time, it
1795 was underthrust as a slab towards the west beneath the Mylonite Zone during convergence
1796 (Fig. 21) (Möller and Andersson, 2018; Möller et al., 2015). The deepest underthrust western
1797 part of the slab reached eclogite-facies conditions corresponding to a depth of c. 70 km (1.65–
1798 1.9 GPa, 850–900°C) at c. 990 Ma (Fig. 5; Fig. 21 b), while the adjacent part of the slab (now
1799 the internal section) reached high-pressure granulite-facies conditions (1.1 GPa, 850°C; Fig.
1800 21 b) (Möller et al., 2015; Piñán-Llamas et al., 2015; Tual et al., 2017). Preservation of
1801 prograde zoning in garnet in eclogites attests to faster-than-equilibration prograde
1802 metamorphism (Möller, 1998; Tual et al., 2017). In Fig. 21, we propose that this underthrust
1803 crustal slab was pulled down by foundering of the dense subcontinental lithospheric mantle.
1804 Breakoff of the lithospheric mantle slab triggered exhumation after c. 980 Ma (Fig. 22).
1805 Exhumation took place in two steps (Fig. 22). During the first step, the eclogitized
1806 westernmost part of the Eastern Segment was detached from the deepest part of the segment
1807 and exhumed with an overall eastward vergence, as a single and coherent (eclogite-bearing)
1808 ductile nappe to an intermediate depth of c. 35–40 km (1.1 GPa), where it was juxtaposed to
1809 the granulite-facies internal section. This process is interpreted as eastwards extrusion during
1810 overall E-W convergence (Möller and Andersson, 2018; Möller et al., 2015; Piñán-Llamas et al., 2015; Tual et al., 2015). During the second step, the eclogite-bearing nappe and the granulite-facies internal section were exhumed together, also with an overall eastwards convergence. Accurate geochronology of these two steps remains difficult to establish. Breakoff of the lithospheric mantle slab was probably facilitated by partial melting in the crust. East vergent exhumation was clearly lubricated by abundant partial melting. The crystallization of these leucosome melts is dated between c. 978 and 961 Ma inside the eclogite-bearing nappe (Andersson et al., 2002a; Möller et al., 2015) and between c. 976 and 965 Ma, i.e. in a coeval time interval, in the granulite-facies internal section (Andersson et al., 2002a; Hansen et al., 2015; Möller et al., 2007; Piñán-Llamas et al., 2015; Söderlund et al., 2002).

1820 The Blekinge-Dalarna dolerite dyke swarm intruded in the upper crust in the foreland of the orogen between c. 978 and 946 Ma (Gong et al., 2018; Ripa and Stephens, 2020d; Söderlund et al., 2005). This mafic magmatism is characterized by a within-plate geochemical signature and supra-chondritic Hf isotopic signature (+1 < εHf < +5). It is evidence for asthenosphere upwelling and decompression melting under the cratonic lithosphere of the foreland (Gong et al., 2018; Ripa and Stephens, 2020d; Söderlund et al., 2005). The upwelling may represent a dynamic response in the asthenosphere of the breakoff and foundering of the lithospheric mantle slab under the Eastern Segment at and after c. 980 Ma (Fig. 22).

1828 Eastwards thrusting in the frontal wedge and along the Sveconorwegian front represents a final spasm of the orogeny. Both ⁴⁰Ar/³⁹Ar data and the observation that Blekinge-Dalarna dolerite dykes are sheared in the frontal wedge suggest that it took place after c. 945 Ma (Andréasson and Dallmeyer, 1995; Page et al., 1996a; Stephens and Wahlgren, 2020a; Ulmius et al., 2018).
1833 **6.3.2 Sustained orogenic plateau west of the Mylonite Zone**

1834 Little tectonic activity or metamorphism is dated in the central part of the orogen (western
1835 part of the Idefjorden lithotectonic unit, Kongsberg and Bamble lithotectonic units and eastern
1836 part of the Telemarkia lithotectonic unit) between c. 980 and 930 Ma, suggesting that this part
1837 of the orogen behaved passively during this time interval. In the west, in the orthopyroxene
1838 zone of Rogaland, a scatter of zircon rim U–Pb ages between 1045 and 955 Ma is interpreted
1839 as evidence for protracted high-grade metamorphism with melt-present conditions in the
1840 middle crust (0.45–0.55 GPa) (Fig. 22) (Blereau et al., 2017; Laurent et al., 2018a; Slagstad et
1841 al., 2018). In Fig. 22, we propose that the orogenic plateau formed between 1050 and 1000
1842 Ma west of the Mylonite Zone was sustained throughout the orogeny to c. 930 Ma.

1843

<figure><img src="image_36.png" /></figure>
1845 **Figure 22.** Schematic geodynamic model for the 980–940 Ma time interval. (a) Sketch map

1846 of the Sveconorwegian orogen, with distribution of metamorphism and magmatism. (b)

1847 Interpretative E-W cross section. WGR: Western Gneiss Region. Explanations and references

1848 in the text.

1849

<figure><img src="image_37.png" /></figure>

1851 **Figure 23.** Geodynamic model for the 940–900 Ma time interval. (a) Sketch map of the

1852 Sveconorwegian orogen, with distribution of metamorphism and magmatism. (b) Pressure-

1853 temperature diagram. (c, d) Interpretative E-W cross sections. Explanations and references in

1854 the text.
1855

1856

1857 6.3.3 Late-Sveconorwegian magmatism and associated metamorphism

1858 The volume of late-Sveconorwegian magmatism increases dramatically westwards in the

1859 orogen (Fig. 5; Fig. 7; Fig. 22; Fig. 23). In the Eastern Segment, minor pegmatite and granite

1860 bodies formed between c. 961 and 935 Ma during regional cooling (Möller et al., 2007;

1861 Söderlund et al., 2008b). In the Idefjorden lithotectonic unit, the large biotite + muscovite-

1862 bearing Flå and Bohus granite plutons (c. 932–922 Ma) (Eliasson et al., 2003; Eliasson and

1863 Schöberg, 1991; Lamminen et al., 2011) carry a distinctly peraluminous signature (S-type)

1864 and sub-chondritic epsilon Nd values (-8.4 < ε$_{Nd}$i < -2.7) (Andersen et al., 2001). These

1865 properties imply a metasedimentary source, most probably in the hosting Stora Le-Marstrand

1866 complex (-8.0 < ε$_{Nd930}$ < -3.0) (Åhäll and Daly, 1989). In the Telemarkia lithotectonic unit,

1867 large plutons of the HBG granite suite emplaced between c. 985 and 926 Ma (Fig. 22; Fig.

1868 23). There is a significant spread in geochemical and isotopic composition, reflecting a

1869 diversity of sources and petrogenesis (Andersen et al., 2001; Granseth et al., 2020; Vander

1870 Auwera et al., 2011). The HBG suite is characterized by a distinctly ferroan geochemical

1871 signature (Fig. 14). Experimental petrology and geochemical modelling of the representative

1872 Lyngdal pluton in Vest Agder (Bogaerts et al., 2006; Vander Auwera et al., 2008), suggest

1873 that this granodiorite crystallized at shallow conditions corresponding to pressures between

1874 0.2 and 0.4 GPa from a wet (5–6 wt% H$_{2}$O), oxydized (QFM +1) and hot (c. 975 °C) magma.

1875 This magma can be generated by partial melting of an amphibole-rich mafic source (with c.

1876 1.5 wt% H$_{2}$O). The near-chondritic to sub-chondritic Nd isotopic signature of the HBG

1877 plutons (-6.4 < ε$_{Nd}$i < +1.9, n = 7; -2.0 < ε$_{Hfi}$ < +1.7, n = 10) overlaps with the evolution trend

1878 of the crust generated at 1520-1480 Ma (Fig. 8; Fig. 13). This implies sources isotopically

1879 similar to those of the Sirdal magmatic belt. However, the geographical overlap between the
1880 HBG suite and the Sirdal magmatic belt suggests that more refractory lower crustal sources were exploited at higher temperature after 985 Ma for the HBG suite (Granseth et al., 2020; Vander Auwera et al., 2008).

1883 To the west of the Sirdal magmatic belt, the AMC suite formed between 932 and 915 Ma in a crust previously metamorphosed to granulite conditions (1030–1005 Ma). The AMC suite is ferroan and alkalic (Fig. 14). Constraints from experimental petrology indicate that the high alumina basalt parental to the anorthosite plutons is characterized by a too low Mg# (molar Mg/(Mg+Fe) = 0.52) and crystallizes too sodic plagioclase (An55) to be generated by melting of a mantle peridotite (HLCA and TJ compositions; Duchesne et al., 1999; Longhi, 2005; Longhi et al., 1999). Instead, its composition is situated on the thermal divide of the plagioclase + pyroxene liquidus surface at 1.0 to 1.3 GPa, imposing that it was produced by partial melting of a gabbronoritic source (Longhi, 2005; Longhi et al., 1999). Experiments show that compositionally adequate melts in equilibrium with plagioclase and orthopyroxene are found in a temperature range between c. 1180 and 1250 °C at c. 1.1 GPa (Fram and Longhi, 1992; Longhi et al., 1999; Vander Auwera and Longhi, 1994). The Sm–Nd isochron of 1041 ± 17 Ma defined by the high-aluminium orthopyroxene megacrysts (HAOM) hosted in the anorthosite plutons (Bybee et al., 2014) suggests that these megacrysts are restitic crystals from a lower crustal source (Vander Auwera et al., 2014b). The isochron implies that the gabbronoritic source formed at c. 1040 Ma as an underplate (1.1 GPa) and was remelted at c. 930 Ma to form the parental magmas of the AMC suite (Vander Auwera et al., 2014b).
1900 Isotopically, this two stage model is realistic, with overlapping positive epsilon Nd values for the megacrysts (+3.1 < εNd(930 Ma) < +5.9) and mafic rocks in the AMC suite (εNd(930 Ma) < +5.8) (Fig. 13). However, the wide range of Nd isotopic composition of differentiated rocks in the AMC suite (-2.8 < εNd(930 Ma) < +5.8), implies a variety of lower crustal sources and crustal contaminants in the suite, all of them characterized by low water content (Barling et al., 2000;
1905 Duchesne and Wilmart, 1997). This two-stage model requires that hot asthenosphere was
1906 upwelling just under the crust around 930 Ma in order to extensively remelt (high degree of
1907 partial melting) gabbronoritic layers of the 1040 Ma underplate. Heat transfer from the mantle
1908 and the anorthosite plutons to the crust resulted in a second phase of ultra-high temperature
1909 granulite-facies metamorphism (M2; 0.35-0.5 GPa, 900-950 °C, c. 930 Ma) (Laurent et al.,
1910 2018b). Using a conservative geothermal gradient of 20 °C/km, extrapolation of a 900 °C
1911 temperature at 0.4 GPa (15 km depth; M2) to the base of the crust at 1.1 GPa (41.5 km depth)
1912 indeed yield a temperature of 1420 °C, compatible with the presence of asthenosphere at the
1913 base of the crust at c. 930 Ma.

1914

<figure><img src="image_38.png" /></figure>

1916 **Figure 24.** Erosion of the Sveconorwegian orogen. (a) Relative probability diagrams of detrital zircons in present day river sediments and dunes. Peaks in this distribution are attributed to formation of supercontinents Nuna, Rodinia and Gondwana. Compilation of Campbell and Allen (2008). (b) Relative probability diagrams of detrital zircons in Neoproterozoic clastic sediments in the North Atlantic realm, deposited in marine and
1921 continental environment mainly during the Tonian and Cryogenian. The compilation includes

1922 3730 detrital zircons from the Moine Supergroup in Scotland (Kirkland et al., 2008b),

1923 Caledonian Lower and Middle Allochthons in Norway and Sweden (Be'eri-Shlevin et al.,

1924 2011; Bingen et al., 2011; Gee et al., 2015; Kirkland et al., 2007, 2008a; Lamminen et al.,

1925 2015; Zhang et al., 2015, 2016), Timanides in N Norway (Zhang et al., 2015), Northwestern

1926 terrane in Svalbard (Pettersson et al., 2009) and Eleonore Bay Supergroup in E Greenland

1927 (Sláma et al., 2011). Only analyses with discordance < 5% are selected; the ²⁰⁶Pb/²³⁸U age is
1928 selected for zircon younger than 1500 Ma and the ²⁰⁶Pb/²⁰⁷Pb age for older zircons. (c)

1929 Relative probability diagram for magmatic events in the entire Sveconorwegian orogen. The

1930 time intervals for continental lithosphere generation and orogenic plateau development are
1931 highlighted. The similitude in the age distribution between magmatic events in the

1932 Sveconorwegian orogen and the Neoproterozoic strata argues for sourcing in the

1933 Sveconorwegian orogen for these sediments and important transport of detritus northwards
1934 and westwards.

1935 -----------------

1936 **6.4 Post 920 Ma: late- to post-Sveconorwegian collapse and sedimentation**

1937 The orogenic plateau developed during the Sveconorwegian orogeny (Fig. 17 to Fig. 23)

1938 could not be sustained when convergence came to a halt sometime after c. 930 Ma, and it

1939 collapsed. As noted earlier, syn- to late-Sveconorwegian plutons (1066-920 Ma) exposed

1940 today define a rather uniform depth of intrusion of c. 16 km (0.4-0.5 GPa, Table 1) (Charlier

1941 et al., 2010; Coint et al., 2015; Eliasson et al., 2003; Vander Auwera et al., 2014a; Vander

1942 Auwera and Longhi, 1994). Removal of this c. 16 km thick overburden took place by a

1943 combination of late- to post-Sveconorwegian erosion and extensional tectonics.

1944 Mapping and characterization of extensional shear zones that could explain exhumation of

1945 amphibolite-facies gneiss complexes (infrastructure of the orogenic plateau) relative to low-
1946 grade supracrustal rocks (superstructure) are still in their infancy (Persson-Nilsson and

1947 Lundqvist, 2014; Torgersen et al., 2018; Viola et al., 2011). Extensional reactivation of the

1948 main shear zones in the orogen, including the Sveconorwegian Front, the Mylonite Zone and

1949 the Bamble-Telemarkia boundary zone, are documented between c. 930 and 860 Ma by

1950 muscovite and biotite ⁴⁰Ar/³⁹Ar data (Andréasson and Dallmeyer, 1995; Mulch et al., 2005;

1951 Page et al., 1996b; Viola et al., 2011). Extension was associated with regional cooling, as

1952 documented by regional scale titanite U-Pb and amphibole, muscovite and biotite ⁴⁰Ar/³⁹Ar

1953 data (Bingen et al., 1998; Connelly et al., 1996; Johansson et al., 2001; Page et al., 1996a;

1954 Page et al., 1996b; Söderlund et al., 1999; Ulmius et al., 2018; Verschure et al., 1980; Wang

1955 et al., 1998).

1956 Pegmatite bodies represent the youngest magmatism of regional significance in the orogen

1957 between c. 914 and 900 Ma (Hetherington and Harlov, 2008; Müller et al., 2015; Müller et al.,

1958 2017; Pasteels et al., 1979; Scherer et al., 2001; Seydoux-Guillaume et al., 2012). Pegmatites

1959 are locally abundant in the gneiss complexes of the Telemarkia and Idefjorden lithotectonic

1960 units. In the Telemarkia lithotectonic unit, they formed shortly after the regional scale titanite

1961 U-Pb age of c. 918 Ma, interpreted to record regional cooling below c. 600 °C. The

1962 pegmatites are not genetically related to any exposed granite pluton and therefore they

1963 represent small individual batches of fluid-rich melt sourced locally in the gneiss complexes

1964 (Müller et al., 2015; Müller et al., 2017). Their relation to the extensional collapse of the

1965 orogen and the source(s) of fluids necessary to generate the fluid-rich melts remain enigmatic.

1966 Absence of widespread late-Sveconorwegian sedimentation inside the orogen is suggested

1967 by a lack of post-Sveconorwegian sedimentary cover below the sub-Cambrian peneplain in

1968 southern Norway and Sweden (Gabrielsen et al., 2015). The Neoproterozoic Visingsö Group

1969 deposited between c. 885 and 740 Ma along the Sveconorwegian Front is directly and

1970 unconformably overlying Paleoproterozoic basement (Loron and Moczydłowska, 2018;
1971 Moczydłowska et al., 2018; Pulsipher and Dehler, 2019). These observations suggest that the Sveconorwegian orogen was erodible, i.e. above sea level, at the end of the Sveconorwegian orogeny.

1974 The rock record supportive for clastic transport towards the Fennoscandia foreland is only local (Almesåkra Group) (Ripa and Stephens, 2020d). However, thermochronological data suggest heating of the foreland up to c. 220 °C some 150 km east of the Sveconorwegian Front between 944 Ma and 851 Ma, which corresponds to burial of the present day surface to c. 7 km (Guenthner et al., 2017). Clastic transport towards the north and west of the Sveconorwegian orogen is, instead, quite well established, as Neoproterozoic continental and marine sediments abound in Caledonian parautochthons and allochthons, Svalbard, and the Timanides (Nystuen et al., 2008). These sediment sequences contain detrital zircons with age between 1700 and 1500 Ma and 1280 and 900 Ma (Bingen et al., 2011; Gee et al., 2015; Kirkland et al., 2007; Kirkland et al., 2008b; Pettersson et al., 2009; Sláma et al., 2011; Strachan et al., 1995; Strachan et al., 2013; Zhang et al., 2015). These sequences can be, at least partly, sourced directly from within the Sveconorwegian orogen (Fig. 24).

1986 -----------------
<figure><img src="image_39.png" /><figcaption>**Figure 25.** Summary matrix of Sveconorwegian orogeny. The first line provides the time slices of events. Line 2 represents the lithotectonic units colour coded following the inset map at the bottom. Line 3 (Tectonics) indicates if the tectonic regime is compressional or extensional in the specific time interval and geographic location. Note that for the time interval between 1280 and 1080 Ma, the tectonic regime is extensional in Telemarkia while it is compressional in the Bamble and Kongsberg Lithotectonic Units. Lines 4-5 (Pressure-Temperature) provides the conditions of peak metamorphism and line 6 (P-T path) indicates where this metamorphism follows a clockwise P-T path. Line 7 (magmatism) records plutonic events, with color coding for felsic (purple), mafic (green) or anorthosite (brown). Line 9 records known sedimentation in intermontane basins or foreland basins. The bottom-right figure is an interpretative transect in map view through the orogen between the Amazonia hinterland and Fennoscandia foreland, showing the three step growth of the orogenic plateau from the centre of the orogen. This three steps growth is associated with three steps of retro-delamination of the sub-continental lithospheric mantle (SCLM) towards the foreland, as</figcaption></figure>
2002 recorded by metamorphism with a high-pressure signature. The western part of the orogen is
2003 characterized by upwelling and shallow asthenosphere during the entire orogenic period.

2004 Explanations in the text.

2005 -----------------

2006 **7 Conclusions**

2007 Published models attempting to explain the tectonic evolution of the Sveconorwegian orogeny
2008 vary widely even with respect to their first-order features and boundary conditions. They
2009 range from end-members involving continental collision between Fennoscandia and another
2010 continent to accretion in the absence of collision (Fig. 3). This variation reflects the difficulty
2011 to translate the observed geological record and the analytical data into geodynamic processes
2012 in the Proterozoic.

2013 Based on a review of data and concepts, we favour a model of large, hot and long-duration
2014 continental collision at the margin of Fennoscandia between c. 1065 and 920 Ma, as
2015 synthesized in Fig. 17 to Fig. 23, and by a matrix in Fig. 25. The plate tectonic setting of the
2016 pre-collision events, between 1280 and 1080 Ma, remains uncertain, although it was possibly
2017 a continental back arc setting (Fig. 17; Fig. 18). Although not strictly necessary, the model is
2018 adjusted into a classical Rodinia assembly framework, involving a Baltica–Laurentia–
2019 Amazonia collision (Fig. 1; Fig. 8, Fig. 12; Fig. 13; Fig. 16).

2020 The width of the orogenic zone, the evidence for protracted and widespread crustal melting
2021 and high-temperature metamorphism reaching UHT conditions, the evidence for high pressure
2022 metamorphism recording crustal thickening, the growth of the orogenic zone towards the
2023 foreland, the juxtaposition of low-grade supracrustal rocks and high-grade gneiss complexes,
2024 and the lack of syn-orogenic marine sedimentary sequence, argue for a collisional orogeny.
2025 We suggest that an orogenic plateau started to form around 1280 Ma in the Telemarkia
2026 lithotectonic unit, first as a Cordillera-style (back-arc) orogenic plateau, and that it grew
2027 stepwise both towards the hinterland and foreland, as a Tibetan-style (collisional) orogenic
2028 plateau. Shallow asthenosphere conditions were maintained in the western part of the
2029 orogenic belt at least up to c. 930 Ma, when the formation of anorthosite plutons took place
2030 by remelting of mafic underplates themselves formed at c. 1040 Ma. Formation of the
2031 orogenic plateau was paired with retro-delamination and foundering of the sub-continental
2032 lithospheric mantle. This process is recorded by compression and regional metamorphism
2033 with an increasingly higher pressure signature towards the foreland followed by exhumation.

2034 Three stages of lithosphere foundering are inferred, one at c. 1150-1120 Ma under the
2035 Bamble and Kongsberg Lithotectonic units, one at c. 1050 Ma under the Idefjorden
2036 lithotectonic unit and one at c. 990 Ma under the Eastern Segment. In the Eastern Segment,
2037 peak conditions reached eclogite facies conditions (1.8 GPa-870 °C) and exhumation of
2038 eclogite-bearing units was aided by extrusion of a ductile nappe lubricated by partial melting,
2039 within an overall compressional setting. The increasing peak pressure recorded in time and
2040 space reflects increasing mechanical coupling between the lower crust and colder lithospheric
2041 mantle, as the delamination process progressed toward the Fennoscandia craton.

2042 After c. 930 Ma, convergence came to a halt, the orogenic plateau collapsed, and 16 km of
2043 overburden was removed by extension and erosion.

2044 **8 Declaration of interests**

2045 The authors declare that they have no known competing financial interests or personal
2046 relationships that could have appeared to influence the work reported in this paper.

2047 **9 Acknowledgements**

2048 The Geological Survey of Norway provided project time to perform this research. T.
2049 Horscroft and M. Santosh are thanked for handling this publication. Mike Stephens and
2050 Aphrodite Indares provided very constructive review of the manuscript.
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