Paleozoic

The metamorphic history of the Precambrian basement need not concern us here. Precambrian basement rocks were, however, at the surface in the Cambrian when they were eroded to form a very flat surface, the Sub-Cambrian Peneplain (Lidmar-Bergström et al. 2013; Gabrielsen et al. 2015). This surface was transgressed during the Cambrian–Ordovician and buried below a sequence of marine mudstones and sandstones. The Caledonian Orogeny had a long, complex history between 500 and 410 Ma, culminating in the collision between Laurentia and Baltica between 420 and 410 Ma when a succession of nappes were thrust eastwards over the basement and some of the Cambrian–Ordovician sediments. The lowermost thrusts are found at the base of and within the sediments that acted as a décollement, and that were metamorphosed to greenschist facies (the so-called phyllites).

The Caledonian mountains collapsed in the early Devonian on extensional faults both within and to the NW of the preserved Caledonian metamorphic rocks. The amount of extension was large, shown by the presence of eclogite-facies rocks, brought up from depths of 60–100 km, in contact with overlying Devonian sediments in west Norway. The Cambrian–Silurian sediments within the Oslo Graben appear to be unmetamorphosed equivalents of the phyllites, and remnants of similar sediments are found scattered over much of southern Sweden, showing that the transgression over the Sub-Cambrian Peneplain was very widespread.

Further clues to the late Paleozoic history of vertical motion of the study area may be derived from the sedimentary basins offshore Norway and in Greenland. By mid-Permian times, the Caledonian mountains in Greenland had collapsed, and the sea transgressed a flat peneplain (Haller 1971). To the west and south of Norway, a similar development took place at that time. Marine Zechstein (upper Permian) sediments lie above the eroded remnants of the Caledonian mountains in the northern North Sea (between southern Norway and Scotland), above subsided Caledonian rocks in the southern North Sea and above remnants of the Variscan mountains in northern Germany (Ziegler 1985; Glennie et al. 2003; Peryt et al. 2010). It is therefore possible that the low-lying, late Permian landscape also included what is now the Southern Scandes.

Mesozoic

Three phases of rifting affected the area to the west of southern Norway, in the Permo-Triassic (c. 250 Ma), Late Jurassic (c. 165 Ma) and Early Cretaceous (c. 110 Ma). Each was followed by a period of continued sedimentation as the basins subsided thermally.

Along much of the coast of southwestern Norway, Jurassic sediments rest on basement, and further offshore they are progressively more deeply buried below Cretaceous and younger sediments (Figs 1a and 2). Three outliers of Jurassic sediments occur along the west coast of Norway (Bøe et al. 2010). Of particular interest for this study is the Upper Jurassic (early to middle Oxfordian), marine Bjorøy Formation encountered near Bergen (Fossen et al. 1997), similar to much of the offshore region, just west of the study area, where Jurassic sediments rest on basement (Fig. 2). Andersen et al. (1999) suggested that several of the Devonian extensional faults within the Caledonian nappes were reactivated as extensional faults during the Permian and Mesozoic, and this view was recently supported by Ksienzyk et al. (2016), who showed that faults in the Bergen area were active at 300–280, 210–190 and 120 Ma. Such faulting is consistent with the former presence of late Paleozoic and/or Mesozoic basins above the rocks that today form the Southern Scandes.

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Fig. 2.

Seismic profile off SW Norway. Noteworthy features are the irregular basement topography (hilly relief, see enlargment) covered by Upper Jurassic sediments (see Fig. 4b) and the pronounced tilting and truncation of Neogene and older strata towards the present-day landmass below Quaternary strata, a configuration that is indicative of late uplift, as seen offshore many elevated passive continental margins around the world (Japsen et al. 2012a). The base of the Neogene strata also represents a pronounced low-angular unconformity representing removal of c. 750 m of section (Japsen et al. 2010). Data courtesy of Statoil; seismic interpretation by L. N. Jensen, Statoil and C. Andersen, GEUS.

In Late Cretaceous–Danian time, pelagic ooze accumulated at the bottom of the epicontinental sea that covered much of NW Europe. Today these deposits, which contain very limited amounts of siliciclastic material, form the main constituent of the Chalk Group (or Shetland Group), which reaches thicknesses greater than 1 km along the coast of Norway south of 59°N but is totally eroded below the base-Quaternary unconformity towards the landmass (Scholle 1977; Ziegler 1990; Japsen 1998). Sømme et al. (2013b) found that the Chalk Group in this region contains no mappable siliciclastic units and that there is no evidence of progradational units originating from a hinterland to the east. Sømme et al. (2013b) combined these observations with the volume of sediment stored in point-sourced depocentres along the margin to conclude that the relief of what is now onshore southernmost Norway was less than 500 m in the Late Cretaceous.

Cenozoic

Prior to the onset of seafloor spreading in the NE Atlantic at 55 Ma, there appears to have been minor uplift and erosion of a generally low-lying Norwegian area, shown by the appearance of small sand fans off western Norway during the Paleocene (Ahmadi et al. 2003). Uplift of Norway at this time must, however, have been much less than that of Scotland, as most of the sand forming the Paleocene fans in the Viking and Central Graben areas came from the Moray Firth area and the Orkney–Shetland ridge (Ahmadi et al. 2003). Sømme et al. (2013b) estimated the Paleocene palaeotopography in southernmost Norway as less than 500 m.

Hemipelagic, deep-marine sediments of Eocene age are present offshore along the coast of southern Norway (Heilmann-Clausen et al. 1985; Faleide et al. 2002). In fact, a compilation of stratigraphic and sedimentological evidence from NW Europe led Knox et al. (2010) to suggest the presence of an Eocene Baltic Seaway from western Siberia to the North Sea, across southern Scandinavia. In that case, only the present-day highest summits in southern Norway would have emerged from the Eocene seas. In agreement with this conclusion, Sømme et al. (2013a) estimated the maximum early Eocene palaeotopography of southwestern Norway to be only c. 200 m.

During the early Oligocene, major clastic wedges prograded southwards from present-day Norway and reached their maximum extent at 29 Ma (Schiøler et al. 2007). This change in depositional environment from Eocene times indicates a phase of uplift and erosion of southern Norway that started at about 33 Ma, a few million years before the maximum extent of the Oligocene prograding wedges (Gregersen et al. 1998; Clausen et al. 1999; Faleide et al. 2002; Jarsve et al. 2015). Jarsve et al. (2015) showed that the creation of accommodation space during the Oligocene subsidence was out of phase relative to eustatic sea-level changes, and was controlled mainly by regional-scale differential vertical movements; that is, uplift and erosion of the hinterland landmasses (southern Norway) occurred concurrently with basin subsidence. Sømme et al. (2013a) estimated the maximum early Oligocene palaeotopography of southwestern Norway to reach elevations above 1.6 km north of Sognefjord (north of 61°N).

The Oligocene–Miocene transition was a period of distinct change in the depositional environment in the eastern North Sea Basin, south and SW of Norway (Rasmussen 2004, 2017). In the late Oligocene, the Danish Basin south of the Oligocene deltas was sediment-starved and hemipelagic sediments accumulated, but in the earliest Miocene, delta systems prograded into the basin from the NNE from present-day Norway and southern Sweden. Three periods of delta progradation took place, but whereas braided river systems dominated the fluvial system in earliest Miocene, meandering rivers became increasingly dominant over time. According to Rasmussen (2014, 2017), early Miocene uplift and erosion in the Scandinavian hinterland resulted in progradation of deltas up to 300 m thick and to formation of the braided fluvial systems and deposition of thick, coarse-grained sequences across much of the present-day Danish onshore area (see also Rasmussen et al. 2010). The occurrence of both immature and mature sediments in these sequences reflects erosion of weathered as well as newly exposed basement and shows that these sequences contain both reworked sediments and erosional products from strongly weathered basement. Early Miocene exhumation also affected the present-day offshore areas west and south of southern Norway (Japsen et al. 2007, 2010) as well as southern Sweden (Japsen et al. 2016).

In the late Miocene, marine sediments reaching a thickness of 200–300 m were deposited in a gently subsiding basin west of present-day southern Norway (Fyfe et al. 2003) in water no more than a few hundred metres deep. However, uplift of mid-Norway and the Northern Scandes in the late Miocene–early Pliocene resulted in deposition of the prograding Molo Formation (Eidvin et al. 2007; Løseth et al. 2017). Major progradation into a marine basin at least a kilometre deep also occurred off southern Norway during the Pliocene (Fyfe et al. 2003), indicating both significant uplift of the source area onshore and subsidence of the basin to provide the accommodation space. Japsen et al. (2007) used palaeothermal and palaeoburial data to show that a final phase of exhumation of the Danish Basin began in the early Pliocene.

In summary, geological evidence defines an episodic history of burial and exhumation during the Phanerozoic development of southernmost Norway and surrounding regions. Thick covers of Paleozoic sediment and Caledonian nappes accumulated over the Precambrian basement, but across most of the area, the Precambrian basement was subsequently exhumed; where Jurassic sediments rest on Precambrian basement, the exhumation occurred prior to Jurassic reburial. Subsequently, three main phases of uplift affected southern Norway after the onset of seafloor spreading in the NE Atlantic: at the Eocene–Oligocene transition at c. 33 Ma, in the early Miocene at c. 23 Ma and in the early Pliocene at c. 4 Ma (Stoker et al. 2005; Japsen et al. 2007, 2010).

Elevated plateaux as evidence of Cenozoic uplift

Early geomorphologists regarded the combination of elevated plateaux and deeply incised valleys of the Southern Scandes to be the result of Tertiary uplift. Reusch (1901) therefore made a clear distinction between the young, pre-glacial (fluvial) incised valleys and the near-flat, elevated plateaux, for which he coined the term palaeic (old) relief (e.g. Hardangervidda at 1.2 km above sea-level (a.s.l.); Figs 3 and 4). Wråk (1908) identified four stepped surfaces in southern Norway, which could be followed eastwards along the flanks of the major valleys at high positions in the landscape. He argued that these steps had formed as a result of pronounced uplift events along the western margin of Norway, and that some of these valleys had later been captured by west-flowing rivers. Ahlmann (1919) had a similar view to Wråk (1908) and regarded the ‘viddas’ (low-relief surfaces at high elevation) of southern Norway as representing uplifted peneplains. He also pointed out that the earlier water divide had a more westerly position than at present. Gjessing (1967) was inspired by the idea of Büdel (1963), who argued that deep weathering was important for the evolution of landscapes. Gjessing (1967) thus described the Palaeic Relief as ‘hills and basins interconnected with passes’ and showed that deep weathering and subaerial slope processes were important for the shape of the palaeic forms. He thus regarded the landscape as ‘the paleic surface’ (i.e. one surface and unrelated to base level), and thought that the landforms were formed during warm climates and not by tectonic events.

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Fig. 3.

Topographic profile across the western flank of the Southern Scandes, from the coast to Hardangervidda (location is shown in Fig. 1a and c). Three lines are shown on the topographic profile; the actual profile is represented by the middle line, and the maximum and minimum topographies within a swath 10 km on each side of the profile are shown as the upper and lower lines, respectively. The flank dips steadily from elevations around 1000 m a.s.l. to the coast, but between the deep valleys that cut into the flank, the landscape is dominated by irregular hills cut by linear valleys; the ‘hilly relief’ and ‘fracture valleys’ (Rudberg 1960; Lidmar-Bergström 1995), characteristic features of Mesozoic weathering during a warm and humid climate. The presence of an Upper Jurassic outlier near Bergen confirms the Mesozoic age of the sloping basement surface that is characterized by a hilly relief to high elevations (Fig. 1c), and its thermal maturity (VR) shows that it has been buried by at least 1 km of Jurassic and younger sediments. This Mesozoic peneplain or denudation surface continues offshore, west of Bergen, where it is preserved below a cover of Jurassic and younger rocks (Fig. 2). At elevations of c. 1200 m a.s.l., the profile is dominated by the subhorizontal Hardangervidda, which is a low-relief plateau formed by post-Caledonian erosion (Fig. 4a) and the third-lowest of the four steps (Level III) in the Palaeic Relief as defined by Lidmar-Bergström et al. (2000). The profile thus reveals the two main components of the landscape of the southernmost Scandes: a tilted Mesozoic peneplain (green dashed line) with a hilly relief along the SW coast, cut by deep valleys and truncated by the subhorizontal, Hardangervidda surface (blue dashed line). Comparison with the similar configuration of surfaces defining the South Swedish Dome suggests that Hardangervidda is a Cenozoic erosion surface (Lidmar-Bergström et al. 2013) and our AFTA data show that it formed during the Miocene.

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Fig. 4.

Contrasting landscape elements of the Southern Scandes. (a) Hardangervidda. (a1) View towards the west from the northwestern part of Hardangervidda at c. 1150 m a.s.l. (photo location is shown in (a2)). The rolling topography of the low-relief surface at 1100–1200 m a.s.l. is cut by deep glacially eroded valleys of the Hardangerfjord. (a2) Elevation illustrating the low relative relief of Hardangervidda; detail from Figure 1. (a3) Geological profile across Hardangervidda (location is shown in (a2)). The metamorphosed Cambrian–Ordovician sediments (phyllites) rest in place on the Sub-Cambrian Peneplain where the basal Caledonian thrusts are within the phyllites. Elsewhere, the basal Caledonian thrust forms the contact between the phyllites and basement (Schiphull 1974). The phyllites (and thus the Sub-Cambrian Peneplain) are preserved below Hardangervidda in several synclinal folds, and this demonstrates that Hardangervidda is a post-Caledonian erosion surface. (b) Remnants of sub-Mesozoic hilly relief and fracture valleys (‘sprekkedale’). The photographs show areas between 800 and 1100 m a.s.l. Photo locations are shown in (b4). (b1), (b2) Hilly relief at elevations around 1000 m (see also Green et al. 2013, fig. 4). The hilly relief has been modified by overriding ice, resulting in enhanced relief (see Olvmo et al. 1999); some glacial features have formed, such as a small cirque on the flank of the left-hand hill in (b1) (Kvernafjellet, 965 m a.s.l., rises 190 m above the lake in the foreground), and there are ubiquitous minor forms such as striae and shatter marks. (b3) Detail of a typical, linear fracture valley at an elevation of c. 950 m a.s.l. It should be noted how the fracture pattern controls the side of the valley as well as its direction. (b4) Digital elevation model of the area east of Stavanger shown in Figure 1c. The area is criss-crossed by fractures that have weathered to form fracture valleys and hilly relief structures, characteristic features of sub-Mesozoic denudation surfaces elsewhere in Scandinavia. The deepest of the valleys (e.g. Lysefjord) have been incised to their present depths by large valley glaciers, but the entire area has been modified by glacial erosion to some extent. Nonetheless, pre-glacial landscape elements occur throughout the landscape, in agreement with the conclusion of Lidmar-Bergström et al. (2013) that the flanks of the Southern Scandes are remnants of a sub-Mesozoic peneplain.

Riis (1996) correlated offshore geology and onshore morphological elements in Scandinavia and interpreted deep weathering profiles encountered in the mountainous areas to be relics of Mesozoic denudation. From this, Riis (1996) inferred that the enveloping summit level originated as a peneplain in the Jurassic. He correlated the main Palaeic Relief of the Southern Scandes with the Tertiary plains with residual hills in northern Sweden and based on offshore sediment volumes he concluded that Scandinavia had been uplifted in two phases, Paleogene and Neogene.

Lidmar-Bergström et al. (2000) used the summit surfaces to identify four main steps with a vertical separation of 1500 m within the Palaeic Relief of the Southern Scandes, and interpreted each of them as the result of different episodes of fluvial incision to a base level following uplift. Lidmar-Bergström et al. (2013) argued that the Palaeic Relief reached its present-day elevation after major late Cenozoic uplift because the incision of the deep and narrow valleys of the Southern Scandes begins at the lowest level of the Palaeic Relief at about 800–1200 m a.s.l. No modern mapping of the surfaces in Norway has been done, except for an area around northern Gudbrandsdalen valley at c. 62°N (Bonow et al. 2003). However, this mapping documented the presence of stepped surfaces at high elevation along the main valleys as well as the fluvial origin of these valleys. The study thus confirmed the results of earlier studies of the Palaeic Relief (e.g. Reusch 1901; Lidmar-Bergström et al. 2000).

Some researchers have suggested that the Palaeic Relief is the result of highly efficient and extensive glacial and periglacial erosion (Steer et al. 2012; Egholm et al. 2017). However, Kleman (2008) in Baffin Island and Hall et al. (2013) in western Norway demonstrated that the plateaux were mainly unaffected during the glaciations owing to cold-based, non-erosive, conditions. Significant erosion took place predominantly in valleys where the ice was warm-based. Recently, Andersen et al. (2018) used cosmogenic isotope data to show that plateau erosion accounts for only 10% of the Quaternary erosion output from the Sognefjorden drainage basin (just north of our study area). Further, Steer et al. (2012) and Egholm et al. (2017) ignored the observation that the glacial valleys exploited pre-existing fluvial valleys (Bonow et al. 2003). However, the preservation of fluvial valley shoulders on the plateaux shows that glacial erosion cannot have removed a 300–400 m thick rock column as suggested by Steer et al. (2012). They also failed to take into account that low-relief plateaux similar to those in the Norwegian landscape occur in both glaciated and non-glaciated areas (Lidmar-Bergström et al. 2000; Bonow et al. 2009; Green et al. 2013); for example, Lidmar-Bergström et al. (2000) showed that major landscape features such as high plateaux, a great escarpment and a coastal plain are similar in southern Norway and eastern Australia. These observations show that the formation of elevated plateaux along passive continental margins is not related to glacial action.

Clayey saprolite, fracture valleys and hilly relief as evidence for Mesozoic weathering

Lidmar-Bergström (1995) showed that deep weathering of the fractures in basement rocks containing feldspars by meteoric water in a warm and humid climate can lead to the formation of distinctive landscape types characterized by undulating hilly relief and by fracture valleys (‘sprekkedale’). The meteoric water penetrated fractures and altered feldspar to kaolinite that was subsequently preferentially eroded, forming valleys typically up to 200 m deep between areas of unweathered rock. The warm and humid conditions necessary for the formation of kaolinitic saprolites prevailed in Scandinavia during the Rhaetian–Campanian (Lidmar-Bergström 1982). The resulting hilly relief with associated saprolites was subsequently preserved below a cover of Mesozoic and younger rocks, as at Ivö in southern Sweden where the hilly relief is capped by Upper Cretaceous limestone and chalk (Lidmar-Bergström 1989). Where these protective sedimentary covers have been removed in the recent past, the hilly relief with remnants of kaolinitic saprolite between the hills has been re-exposed and the surface of the basement hills slightly altered, in particular by glacial action (see Lidmar-Bergström et al. 2017, fig. 3).

There is evidence of late Mesozoic weathering of basement both onshore and offshore southern Norway. Seismic lines around Norway show a pattern corresponding to hilly relief at top basement below Jurassic sediments (Fig. 2), and Riber et al. (2015) showed evidence of subaerially weathered basement rocks on the Utsira High (off SW Norway) prior to burial at the Jurassic–Cretaceous transition. Migón & Lidmar-Bergström (2001) mentioned reports of kaolin at several localities around southern Norway. Lidmar-Bergström & Näslund (2002, fig. 4) showed an extensive area of hilly relief to the east of the Oslo Rift, where Olesen et al. (2007) found that structural weakness zones containing clay minerals, including smectite and kaolinite that, to a large extent, are the result of chemical weathering of pre-existing fracture zones during subtropical conditions in the late Mesozoic. Olesen et al. (2007) argued that the weathering penetrated deep into the fracture zones and was preserved below a sedimentary cover until exhumation removed the cover and most of the weathering products while preserving clay zones at depths of 200–300 m.

The Norwegian strandflat is an uneven and partly submerged rock platform along the coast (Reusch 1894). Holtedahl (1998) thought that the strandflat probably formed in late Pliocene or Pleistocene times as a product of glacial erosion, marine erosion and subaerial weathering, and that it coincides with an exhumed pre-Jurassic surface in some coastal areas. Fredin et al. (2017) obtained ages of 221–206 Ma (Norian, Late Triassic) from K–Ar dating of authigenic, syn-weathering illite from saprolitic remnants at Bømlo on the strandflat of southwestern Norway (Fig. 1a). Assuming the K–Ar dates to be accurate, they imply that a surface close to at least part of the strandflat must have been near the Mesozoic surface at that time. Its preservation is unlikely unless it had been covered by sediments in Mesozoic and Cenozoic times that were stripped off again more recently. The known kaolinitic saprolites on the strandflat of west and south Norway (Migón & Lidmar-Bergström 2001) cannot, however, have formed during the Norian because formation of kaolin requires warm wet conditions, and northern Europe was a desert during the Norian. Warm and wet conditions prevailed from the Rhaetian to Campanian, so kaolin is more likely to have formed during this period (Lidmar-Bergström 1982).

Peneplains and topographic highs of Scandinavia

Stratigraphic landscape analysis provides evidence of uplift and subsidence in shield areas, using cross-cutting relationships between peneplains and stratigraphic constraints from the sedimentary cover that preserved them to construct a relative chronology for surface formation and tectonic events (Lidmar-Bergström et al. 2013). Here we adopt the definition that all surfaces graded to a base level are peneplains, irrespective of their detailed form (hilly or flat) (Green et al. 2013; Lidmar-Bergström et al. 2017). In any case, the formation of a peneplain involves kilometre-scale erosion of rock during valley incision and valley widening by running water and slope processes, whereas the final shape of the peneplain depends on the climatic conditions during its formation.

Lidmar-Bergström et al. (2013) described three topographic highs that characterize Scandinavia: the Northern Scandes, the Southern Scandes and the low South Swedish Dome (2.1, 2.5 and 0.4 km a.s.l., respectively; Fig. 1a). The South Swedish Dome emerges from below Cambrian cover rocks in the north and east and from below Cretaceous cover rocks in the south and west, and is delimited to the SW by the Sorgenfrei–Tornquist Zone. Lidmar-Bergström et al. (2013) emphasized that these geological constraints permitted a detailed stratigraphic landscape analysis of the South Swedish Dome, which provides the key to understanding the Phanerozoic history of vertical movements of Scandinavia.

Three major peneplains define the South Swedish Dome of which two are re-exposed peneplains (Lidmar-Bergström et al. 2017). The sub-Cretaceous peneplain, characterized by undulating hilly relief and showing up to 200 m difference between the summit of basement hills and the fresh basement below thick kaolinized basement rocks (saprolite), stands in stark contrast to the extremely flat Sub-Cambrian Peneplain. Both re-exposed surfaces developed close to former sea-levels, were subsequently transgressed, and were buried below sedimentary covers. The preservation of these peneplains below Paleozoic and Upper Cretaceous cover rocks, respectively, documents that uplift of the land surface had been followed by subsidence and burial, and that uplift and erosion in the recent past has brought the peneplains back to the surface. Cross-cutting relationships between these re-exposed and tilted peneplains and a third, horizontal plain with residual hills (South Småland Peneplain) demonstrate that the latter is younger than the sub-Cretaceous peneplain and is thus of post-Cretaceous age.

Lidmar-Bergström et al. (2013) showed that the three relief types of the South Swedish Dome can be identified all around Scandinavia, and that their presence demonstrates phases of uplift/denudation and subsidence/burial of Scandinavia during the Phanerozoic. They argued that the subhorizontal peneplains of the Northern Scandes, the Southern Scandes and the South Swedish Dome are Cenozoic erosion surfaces because these subhorizontal plains truncate tilted, sub-Mesozoic peneplains (Fig. 3). The flanks were thus at the surface and weathered during the Mesozoic but were subsequently covered and tilted prior to their re-exposure during relatively recent uplift.

Key landscape elements of the southernmost Scandes

Overall, our study area, the southern half of the Southern Scandes, represents half of an elongated dome with slopes dipping to the west, south and SE from high ground in the centre (Figs 1b and 3). A subhorizontal plateau, Hardangervidda, with a very low relative relief, occupies much of the higher ground, at elevations of c. 1200 m. Hardangervidda corresponds mainly to the third-lowest of the four steps (Level III) in the Palaeic Relief as defined by Lidmar-Bergström et al. (2000). Level IV occurs some hundreds of metres below Hardangervidda and corresponds to the bottom level of the Palaeic Relief.

The slopes are incised by deep and narrow valleys below Level IV (Lidmar-Bergström et al. 2000), commonly showing distinct fluvial character, whose final development involved enlargement by valley glaciers (Bonow et al. 2003). Many irregular hills cut by linear valleys also occur between the deep valleys. At low levels around much of the region, we find irregular bedrock surfaces, which Lidmar-Bergström (e.g. Lidmar-Bergström 1995) described as undulating hilly relief and fracture valleys. In coastal regions, an uneven and partly submerged rock platform, the strandflat, occurs (Holtedahl 1998). In places, the amount of glacial erosion has been so great that it has obliterated much of any older landscape elements, but as we discuss below, there appear to be areas both on the top and on the flanks of the dome where substantial pre-Pleistocene landscape elements have been preserved.

Previous thermal history studies

A number of studies have applied apatite fission-track thermochronology in different parts of southern Norway to investigate the exhumation history and the development of the present-day topography of the region. Earliest studies revealed that Precambrian basement rocks at outcrop are characterized by apatite fission-track ages of c. 250 Ma or younger (Andriessen 1990; Grønlie et al. 1994; Rohrman et al. 1995, 1996). These results, together with track-length data, were interpreted as showing that the samples cooled from temperatures of 100°C or above in Triassic–Jurassic times, reflecting Mesozoic exhumation from depths of several kilometres. Olaussen et al. (1994) used onshore apatite fission-track data and offshore seismic data to infer that Early Triassic and ‘mid’ Jurassic episodes of exhumation removed any post-rift basin fill from the Upper Carboniferous–Permian Oslo Rift and the Skagerrak Graben. Rohrman et al. (1995) further interpreted their data as defining a later, Neogene phase of exhumation that formed a domal uplift of southern Norway, responsible for the modern-day topography of the region. Early apatite fission-track studies in southern Sweden reached broadly similar conclusions (Zeck et al. 1988; Larson et al. 1999; Cederbom et al. 2000; Cederbom 2001; Huigen & Andriessen 2004).

Redfield et al. (2004, 2005a,b) confirmed the importance of Mesozoic cooling and revealed major offsets across fault zones along the west coast of central and northern Norway. Data from a number of studies were compiled by Hendriks et al. (2007), demonstrating a consistent regional pattern of variation in apatite fission-track age across the study area (Fig. 5a) and further afield into Sweden and Finland. Youngest ages, typically less than 150 Ma, are found in the centre of the Southern Scandes and further north, somewhat older ages up to c. 250 Ma are found around the coast of southern Norway and into Sweden, and much older ages in central and northern Sweden and Finland. The compiled data show a complex relationship between apatite fission-track age and mean confined track length (Fig. 5b), suggesting significant variation in thermal history across the region.

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Fig. 5.

Apatite fission-track data from this and previous studies (Hendriks et al. 2007; Japsen et al. 2016). (a) Map of apatite fission-track ages in southernmost Norway. (b) Mean track length v. apatite fission-track age for Scandinavian outcrop samples. Data from this study are consistent not only with compiled data from Norway (Hendriks et al. 2007) but also with data from southern Sweden (Japsen et al. 2016), suggesting a coherent thermal history framework across the region. It should be noted that measured fission-track ages from outcrops across southern Norway are dominated by values less than 250 Ma, emphasizing the importance of post-Paleozoic events in controlling the development of the Southern Scandes.

Japsen et al. (2007, 2010) applied AFTA, vitrinite reflectance (VR) and sonic data in wells to study the exhumation history of the offshore sedimentary basins south and west of southern Norway (Fig. 1a). Japsen et al. (2010) integrated their results with evidence from offshore stratigraphy and suggested that southern Norway had undergone three phases of Cenozoic uplift and erosion, starting at c. 33, c. 24 and at c. 4 Ma. Thus, a number of AFT studies have supported earlier inferences from geomorphological investigations that the mountains of southern Norway were uplifted within the relatively recent geological past.

Based on a reinterpretation of a number of apatite fission-track datasets, including Rohrman et al. (1995) and other published and unpublished sources, Nielsen et al. (2009) presented an alternative view, involving more or less continuous cooling/denudation since the Caledonian Orogeny. We discuss this view further in the section on mechanisms.

Later apatite fission-track studies of Norway have been focused around fault movements in the north of the country (Redfield & Osmundsen 2009; Hendriks et al. 2010; Osmundsen & Redfield 2011; Davids et al. 2013). Ksienzyk et al. (2014) provided new apatite fission-track data from southern Norway in the vicinity of Bergen, and also reported the first apatite (U–Th)/He ages from Norway. Their results revealed significant Mesozoic fault offsets, and concluded that the modern-day topography is the result of repeated pre- and post-rift tectonic episodes.

Thus, despite a number of studies over the last 30 years, opinions remain divided in regard to the interpretation of low-temperature thermochronology data and the development of the modern-day topography of southern Norway.

New AFTA data from southernmost Norway

Interpretation

The principles involved in extracting thermal history constraints from AFTA data in studies such as this have been discussed in detail elsewhere (Green & Duddy 2012; Green et al. 2013). The controlling principle is that in samples that are heated to a maximum palaeotemperature and then cooled, the data are dominated by the palaeothermal maximum, and the prior history is overprinted. For this reason, it is not possible to constrain the entire history, so instead we focus on defining the palaeothermal maximum and subsequent palaeothermal peaks, within an overall framework of episodic heating and cooling reflecting the episodic nature of the Phanerozoic development of the study area documented by the geological record and by the contrasting types of bedrock relief as summarized in the previous sections. As explained in detail elsewhere we believe that this type of history is more geologically realistic than those involving slow, monotonic cooling (Green & Duddy 2012; Green et al. 2013, 2018). This belief is based on AFTA data in situations where geological constraints provide key insights into the history, defining not only episodes of cooling/exhumation but also prior periods of heating/burial, which cannot be explained by slow continuous cooling/exhumation histories.

Thermal history solutions, represented by 95% confidence limits on the maximum or peak palaeotemperature and the time at which cooling from that palaeotemperature began, are obtained by comparing measured data with values predicted from candidate thermal histories involving up to three episodes of heating and cooling (three being the maximum number of cooling episodes that can be obtained from AFTA data in most cases). Where three episodes are defined, the earliest episode is defined primarily from the fission-track age data (and represents the episode in which samples cool below c. 110°C and began to retain tracks), whereas the two more recent episodes are defined principally from the distribution of track lengths (see Fig. C.21 of Appendix C in Supplementary material S2). By varying the magnitude of maximum paleotemperature and the onset of cooling in each episode (using assumed heating and cooling rates each of 1°C Ma⁻¹), the range of values giving predictions that are consistent with the measured data within 95% confidence limits is defined.

Thermal history constraints extracted from the AFTA data in each sample, following the strategy outlined above, are summarized in Supplementary Table S1. Figure 6 shows a comparison of the constraints (95% confidence intervals) on the onset of cooling in up to three palaeothermal episodes in each sample; that is, times when samples were hotter than they are today. If we assume that data in individual samples represent synchronous cooling episodes resulting from regional processes such as exhumation or reduction in heat flow, the results in almost all samples from this study can be explained in terms of just three Phanerozoic cooling episodes, beginning in the following intervals (assignment of chronostratigraphic age based on Gradstein et al. 2012): 251–243 Ma; Early to Middle Triassic; 172–164 Ma; Middle Jurassic; 30–21 Ma; Oligocene–early Miocene.

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Fig. 6.

Constraints on the timing of onset of cooling in palaeothermal episodes recognized from AFTA data in this study (Supplementary table S1). Synthesis of all data suggests three dominant episodes of cooling, as indicated by the vertical bands with different colours. In addition, an earlier onset of cooling is seen in three samples, broadly defining a Paleozoic cooling event. The stratigraphic age of Precambrian basement samples plots outside the diagram. Constraints that are not consistent with the preferred timing for the major regional events are highlighted with a yellow background.

It should be noted that these intervals define the uncertainty on the time at which cooling began in each episode, and we do not imply that all cooling took place within these intervals. Three timing constraints that are not consistent with the above synthesis, highlighted in Figure 6, are regarded as statistical outliers. In addition, an earlier cooling episode is seen in three samples, consistent with cooling that began between 411 and 333 Ma, probably reflecting late-stage Caledonian cooling.

Regional synthesis

As we noted earlier, the AFTA data in samples from this study are highly consistent with those from our study of southern Sweden (Japsen et al. 2016), and the three regional cooling events defined above closely correlate with cooling events defined in Sweden; namely, 245–240, 171–142 and 23–15 Ma. It should be noted that the data from southern Sweden defined a late Carboniferous event of cooling and exhumation (beginning between 314 and 307 Ma) that could not be resolved in the data presented here, possibly because of the high palaeotemperatures involved in the Middle Triassic thermal event. Furthermore, the dataset from Sweden also defined a mid-Cretaceous cooling event beginning between 107 and 92 Ma (Albian–Turonian), which reflects the onset of inversion of the Sorgenfrei–Tornquist Zone.

Given the consistency of the two datasets we conclude that results from both areas reflect common events, and that the best estimates of the onset of cooling in these events are as follows:

• 245–243 Ma, Middle Triassic (Anisian);

• 171–164 Ma, Middle Jurassic (Bajocian–Bathonian);

• 23–21 Ma, early Miocene (Aquitanian).

Given the wide timing constraints in individual samples (Fig. 6), somewhat wider uncertainty ranges should probably be applied to the Triassic and Miocene episodes. In the following discussion, we follow this interpretation in terms of three regionally synchronous cooling episodes across southern Scandinavia. On the basis of the regional extent of each of these episodes, we suggest that an explanation in terms of cooling owing to exhumation is the most appropriate explanation for each of these cooling episodes, as also concluded by Japsen et al. (2016) in regard to southern Sweden.

Middle Triassic exhumation leading to removal of Paleozoic cover on the flank of the earliest North Sea rift

Almost all samples cooled from palaeotemperatures above 100°C and began to retain tracks in the Middle Triassic episode at c. 245 Ma. All samples forming the near-vertical profile from Rjukan (located at the southeastern edge of Hardangervidda; Figs 1c and 7a) define Middle Triassic palaeotemperatures >100°C, corresponding to burial below a cover many kilometres thick for any reasonable palaeogeothermal gradient. This cover is likely to have been a thick succession of Caledonian nappes and probably upper Paleozoic–Lower Triassic sediments that covered what is now Hardangervidda prior to the onset of Middle Triassic exhumation. Only three samples, all located close to the west coast, cooled from above 100°C in an earlier episode and these cooled from palaeotemperatures around 90–100°C in the Middle Triassic episode (Fig. 6). Because the rock that was removed was likely to have lithologies similar to the rock that remains, issues connected with possible non-linear thermal profiles discussed below are not significant here. These palaeotemperatures correspond to burial below c. 3 km for a palaeogeothermal gradient of 25°C km⁻¹ or c. 1.5 km for a palaeogeothermal gradient of 50°C km⁻¹ (assuming a palaeo-surface temperature of 20°C).

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Fig. 7.

Palaeotemperature constraints from AFTA v. elevation compared with assumed palaeogeothermal gradients of 25 and 35°C km⁻¹. (a) Constraints from a vertical transect from Rjukan showing palaeotemperatures for the Middle Triassic, Middle Jurassic and early Miocene paleo-thermal episodes. Extrapolation of these constraints along the assumed palaeogeothermal gradient of 25°C km⁻¹ yields palaeotemperatures at present-day sea-level of more than 100°C, about 90°C and about 70°C, respectively. If the early Miocene palaeotemperatures are extrapolated along that gradient they reach the likely palaeo-surface temperature of 20°C at an elevation of 2 km a.s.l. (the projected surface elevation, PSE; Japsen et al. 2014). This implies that prior to the onset of early Miocene cooling, the land surface was well above present-day terrain. (b) Early Miocene paleotemperature constraints for all samples colour-coded by location. The constraints from the Rjukan and Lysefjord profile match a palaeotemperature gradient of 25°C km⁻¹. For this gradient, the projected surface elevation for inland locations is about 2 km a.s.l. (red circle a). For the coastal locations it is slightly lower, about 1.7 km a.s.l. (red circle b). For a palaeotemperature gradient of 35°C km⁻¹, corresponding to measured values for the sedimentary cover offshore (Japsen et al. 2010), the projected surface elevation for coastal locations is about 1.3 km a.s.l. (red circle c). The estimated palaeotemperature for a sample of the Jurassic outlier in Bergen (Fossen et al. 1997) based on a new measurement of vitrinite reflectance (0.38%) is also shown. Coastal locations include Setesdalen.

Rohrman et al. (1995) also reported evidence for Triassic exhumation in southern Norway based on apatite fission-track data, and Johannessen et al. (2013) used apatite fission-track data combined with (U–Th)/He data from the inner Hardangerfjord to define accelerated Permo-Triassic cooling. Fauconnier et al. (2014) found that the Cambro-Silurian sediments below the Jotun Nappe Basal Thrust, just north of our study area, had been heated to temperatures ranging from 320 to 500°C during the main Caledonian burial stage and thus prior to the Middle Triassic exhumation event (Fig. 8). A kilometres-thick cover of Paleozoic–Lower Triassic rocks also accumulated on the Sub-Cambrian Peneplain in southern Sweden prior to Middle Triassic exhumation (Zeck et al. 1988; Japsen et al. 2016). Up to 4 km of Paleozoic strata are preserved in the Baltic Sea and in the Sorgenfrei–Tornquist Zone in the Kattegat (Sigmond 1993; Calner et al. 2013).

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Fig. 8.

Heating and cooling (burial and exhumation) history of basement rocks now exposed from below Upper Jurassic sediments along the coast of SW Norway (samples 361-21, 22, 25). The Sub-Cambrian Peneplain and Hardangervidda are found at a short distance from the coast at elevations around 1 km a.s.l. (Figs 1 and 4). A, The Sub-Cambrian Peneplain that extended across Scandinavia (Lidmar-Bergström et al. 2013) formed after uplift and exhumation that began around 600 Ma in the latest Neoproterozoic (Japsen et al. 2016). B, When the peneplain was fully developed prior to Cambrian transgression, rocks at this surface had cooled to the palaeosurface temperature of c. 20°C. C, Heating to 300°C or more below thick Caledonian nappes (Fauconnier et al. 2014). D, Cooling below palaeotemperatures of 100°C after Middle Triassic exhumation. E, Cooling to near-surface conditions and deep weathering in the Late Triassic (Fredin et al. 2017) after removal of the Paleozoic cover and destruction of the Sub-Cambrian Peneplain. F, Heating to c. 80°C below Upper Triassic to Middle Jurassic sediments. G, Exhumation after Middle Jurassic uplift leading to exposure of the bedrock and formation of a deeply weathered, hilly bedrock surface (peneplain) (Figs 2 and 4b). H, Heating to c. 60°C below Upper Jurassic to Oligocene sediments and preservation of the Mesozoic peneplain. I, Exhumation after early Miocene uplift with samples cooling to a surface temperature of c. 10°C on the strandflat.

The Middle Triassic event is likely to have exhumed Precambrian basement to near-surface conditions in the Late Triassic over parts of our study area. This is supported by K–Ar ages of epigenetic mineral deposits in the Oslo Rift (Ineson et al. 1975) and K–Ar ages of illite from remnants of saprolite on the Utsira High and on the strandflat at Bømlo, SW Norway (Fig. 1c; Fredin et al. 2017).

Kilometre-thick successions of coarse clastic sediments of the Middle–Upper Triassic Skagerrak Formation (Hegre Group) are present west and south of southern Norway and in the Danish Basin (Bertelsen 1980; Goldsmith et al. 2003; www.npd.no). The results presented here and by Japsen et al. (2016) suggest that the Skagerrak Formation includes the erosional products derived from southern Scandinavia after Middle Triassic uplift. Jarsve et al. (2015) likewise concluded that southern Norway was the main source area for the Triassic deposits in the central and eastern North Sea. The Middle Triassic event coincided with the concluding stage of the Oslo Rift, which was characterized by the deep intrusion of minor granites at 249–241 Ma (Larsen et al. 2008).

Middle Jurassic exhumation leading to formation of a Mesozoic peneplain

The effects of the Middle Jurassic cooling and exhumation beginning between 171 and 164 Ma are recorded by most of the samples analysed for this study. Samples at elevations near sea-level cooled from palaeotemperatures around 80–90°C, corresponding to burial below 1.2–2.8 km for palaeogeothermal gradients from 50 to 25°C km⁻¹ (palaeo-surface temperature 20°C). Fredin et al. (2017) showed that the basement in western coastal regions was at the surface in the Late Triassic, so these palaeotemperatures indicate that, along the west coast, a kilometres-thick cover of Upper Triassic–Lower Jurassic sediment must have accumulated onto the basement. As Upper Jurassic sediments rest on basement near Bergen (Fossen et al. 1997) and further west offshore (Fig. 2), the sedimentary cover removed during Middle Jurassic exhumation must thus have been of Late Triassic–Middle Jurassic age. This cover was thus removed prior to re-exposure of the basement and subsequent reburial in Oxfordian time (Fig. 8), thus giving a maximum time of about 14 myr between the onset of Middle Jurassic exhumation and Oxfordian reburial. Because the events we are discussing are of regional extent, it seems likely that this former sedimentary cover extended over a wide area of the present-day onshore prior to Middle Jurassic exhumation.

Jurassic palaeotemperatures in samples from the Rjukan vertical profile (Fig. 7a) are consistent with a palaeogeothermal gradient of 25°C km⁻¹ (although the deepest value seems to be anomalously high probably because the Triassic event is not resolved in this sample, and the single constraint probably represents the unresolved effects of Jurassic and Triassic cooling). A palaeotemperature intercept of 90°C at sea-level implies the presence of a rock column in the Middle Jurassic that reached almost 3 km above present-day sea-level. This cover is likely to have been remnant Cambro-Ordovician metasediments and rocks of the Caledonian Jotun Nappe and any Upper Triassic to Lower Jurassic sedimentary cover that extended this far. Because the Jotun Nappe rocks today extend to an altitude of nearly 2 km a.s.l. just north of Hardangervidda, there may not have been more than a few hundred metres of Mesozoic sediment above the Caledonian rocks in this central part of the study area.

A deeply weathered, hilly bedrock surface (peneplain) formed after the Middle Jurassic denudation event. Today the surface is preserved below a cover of Jurassic and younger sediments in the offshore region (Fig. 2), and it is re-exposed along the western and southern flanks of the Southern Scandes from sea-level to about 1 km altitude, just below Hardangervidda (Figs 1c and 3). We will refer to this surface as a Mesozoic peneplain as it was formed after removal of kilometres-thick rock columns and finally graded to a base level, in this case the sea as documented by the marine, Upper Jurassic sediments in Bergen (Fossen et al. 1997). The peneplain formed by weathering during the warm and humid climate in the late Mesozoic and is thus characterized by hilly relief and fracture valleys. Jurassic and possibly Cretaceous sediments accumulated on the peneplain, which consequently was preserved until its recent exposure. Stratigraphic landscape analysis thus assigns ‘sub-Mesozoic’ as the relative age of the denudation surface, whereas our results document that it is a Mesozoic surface.

Mid-Jurassic uplift affected a wide region in the North Sea (Ziegler 1990), both west and south of the present study area. Underhill & Partington (1993) combined stratigraphic observations with evidence for igneous activity from the central North Sea to conclude that regional, domal uplift resulted from the impingement of a broad (>600 km radius), transient plume head at the base of the lithosphere (‘the central North Sea Dome’). Underhill & Partington (1993) did not, however, consider the continuation of this zone of uplift towards the coast of Norway, which they classified as ‘data poor’. Rohrman et al. (1995), however, presented evidence for major Jurassic exhumation in southern Norway based on apatite fission-track data.

The stratigraphic record of the Danish Basin and the Sorgenfrei–Tornquist Zone contains ample evidence of a pronounced, mid-Jurassic phase of uplift and erosion that began in the Aalenian at c. 175 Ma coeval with pronounced volcanic activity at the southwestern margin of the Baltic Shield (Norling & Bergström 1987; Nielsen 2003; Tappe 2004). The cooling and exhumation resulting from this event is also revealed by AFTA data from southern Sweden and the Danish Basin (Japsen et al. 2007, 2016).

Early Miocene exhumation leading to formation of Hardangervidda

The effects of the early Miocene cooling and exhumation beginning between 23 and 21 Ma are recorded by the AFTA data in all but one sample. Samples close to sea-level cooled below palaeotemperatures around 60–70°C (Fig. 7). This is consistent with a new VR value of 0.38% (see Supplementary material S2) for the Upper Jurassic outlier near Bergen (Fossen et al. 1997). This VR value suggests that the Jurassic sediments were heated to a maximum palaeotemperature of 63°C (Burnham & Sweeney 1989, assuming a heating rate of 1°C Ma⁻¹, consistent with that employed in interpreting the AFTA data) prior to exhumation. The AFTA data also suggest that the most likely timing for the maximum heating was in the Miocene. Such paleotemperatures imply burial by a considerable thickness of younger section, subsequently removed during exhumation.

Results from elevation profiles at Rjukan and other locations in Figure 7 are all consistent with a palaeogeothermal gradient of c. 25°C km⁻¹ in basement rocks for this episode, although coastal locations appear to define a profile that is offset to lower temperatures compared with the inland locations. Linear extrapolation of the palaeotemperature profile from inland locations to a palaeo-surface temperature of 20°C implies that a cover of rocks extended to c. 2 km above present sea-level prior to the onset of Miocene exhumation, whereas the cover above coastal locations reached c. 1.6 km above present sea-level. Whereas rocks from the highest elevations tend to define fairly wide temperature constraints in this event (owing to the low temperatures involved), as shown in Figure 7 extrapolation of the profiles suggest that samples at inland elevations close to that of present-day Hardangervidda cooled from around 40°C. Thus, in this scenario the rock column that was present in the early Miocene must have extended well above all of the present-day landscape within the study area.

Estimating the amount of section that was present above the present-day surface by linear extrapolation of the palaeotemperature profile is not straightforward because the analysis depends critically on the assumption that the palaeogeothermal gradient was linear throughout the entire section when cooling began. This is equivalent to assuming that the missing section and the preserved section were of similar lithologies, such that no significant contrasts in thermal conductivity existed that may have resulted in a non-linear palaeotemperature profile (see fig. 32 of Green et al. 2013). Thermal conductivities measured in rocks of different types show wide variation from around 1 W m⁻¹ K⁻¹ or less to over 5 W m⁻¹ K⁻¹, related to lithology, water content, porosity, temperature and other factors (e.g. Robertson 1988; Clauser & Huenges 1995; Eppelbaum et al. 2014). If the cover rocks that were removed during Miocene exhumation were relatively unconsolidated sediments with lower thermal conductivities compared with the underlying basement rocks, then the palaeogeothermal gradient in the overlying section would have been higher. In that case, the amount of cover required to explain the Miocene palaeotemperatures would be correspondingly reduced, perhaps by a factor of two, based on the range of typical thermal conductivities in the studies cited above.

Therefore we can define two likely end-member scenarios for the thickness of the cover that was present above the level of the present-day Hardangervidda immediately prior to the onset of Miocene exhumation. If the rocks that were eroded during Miocene exhumation were Jotun Nappes and phyllites, then linear extrapolation of the profile in Figure 7 from c. 40°C at 1.2 km a.s.l. suggests that around 0.8 km of rock was present above the level of Hardangervidda. The thickness required could possibly be higher if we apply present-day temperature gradients in the range 15–20°C km⁻¹ as measured in boreholes through Proterozoic–Paleozoic rocks in southern Norway (Slagstad et al. 2009). Alternatively, if the rock that was removed was mostly a Mesozoic to Paleogene sedimentary cover, the amount of rock removed above Hardangervidda could have been as little as 0.4 km. We consider the former scenario is more realistic (Fig. 9), but in either case, a significant thickness of rock must have been removed following the onset of Miocene exhumation, resulting in creation of the planation surface that now forms Hardangervidda.

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Fig. 9.

Schematic illustration of the burial and exhumation history of Southern Scandes along an onshore–offshore profile reaching from the eastern North Sea to Hardangervidda (location shown in Fig. 1). (a) Present-day: the interpretation of the seismic line in Figure 2 is shown depth-converted on the left and a schematic representation of the onshore landscapes (Fig. 4) and geology on the right. Red dashed line: the ‘early Miocene surface’ marks the top of the rock column that was present above the present-day surface prior to early Miocene exhumation estimated from the palaeotemperature constraints from AFTA data onshore (this study) and offshore (Japsen et al. 2010). Green dashed line: the tilted Mesozoic peneplain with a hilly relief below the Jurassic sediments on the seismic line and on the sloping surface SW of Hardangervidda. Blue dashed line: the subhorizontal Hardangervidda surface. (b) Mid- to late Miocene: formation of the Hardangervidda planation surface as part of an extensive peneplain (blue dashed line) after early Miocene uplift leading to erosion by river incision to a base level interpreted as adjacent sea-level. Hardangervidda was then uplifted to its present elevation of c. 1200 m a.s.l. as shown in (a). This profile is produced by removing 800 m from the top of the profile in (c) in agreement with estimates of removed covers owing to early Miocene exhumation for Hardangervidda and for wells offshore (Japsen et al. 2010). (c) Early Miocene: maximum Cenozoic burial following rifting in the Late Jurassic and subsequent post-rift subsidence and burial. This profile is produced by flattening the early Miocene surface in profile in (a) to early Miocene palaeo-sea-level. It should be noted that a rock column of about 800 m covered Hardangervidda, interpreted here as Caledonian rocks, and that a column of Jurassic and younger sediments covered the present-day land surface from the coast to the western margin of Hardangervidda. Jurassic sediments extended at least as far east as the present-day coast (Fossen et al. 1997), and we have chosen to show them as onlapping the basement SW of the Folgefonni Glacier, implying that the etch surface at present-day higher elevations was produced during the Cretaceous. FF, Folgefonni Glacier. AFTA data from Stavanger (ST), LF (Lysefjord) and R (Rjukan).

Where the Mesozoic peneplain is exposed onshore today (Fig. 1c), our results suggest that a protective cover of Upper Jurassic to Oligocene sediment was present prior to the onset of early Miocene exhumation, similar to what is preserved offshore today (Fig. 2). In this case, a significant thermal conductivity contrast is likely between basement and the removed cover, in which case linear extrapolation of the palaeotemperatures to higher elevations, as discussed above, would not be valid. Whereas linear extrapolation of the profile for coastal locations in Figure 7 implies a rock column extending to 1.7 km above present-day sea-level, the arguments above suggest that as little as c. 1 km of overburden may have been present. However, we have adopted a thickness of the sedimentary cover on the coast of 1.3 km, based on a palaeogeothermal gradient of 35°C km⁻¹ through the removed section, which corresponds to the upper limit for present-day values of the geothermal gradient for Mesozoic–Cenozoic sediments in wells in quadrant 17 (Fig. 1a; Japsen et al. 2010).

The early Miocene tectonic phase had great impact on the region. It produced a pronounced regional unconformity in the sedimentary basins offshore the margin of NW Europe (Stoker et al. 2005) and caused the removal of a sedimentary cover of 750–1000 m SW and south of Norway (Fig. 10; Japsen et al. 2007, 2010). In the eastern North Sea Basin, coarse-grained, braided river systems from Norway reached western Denmark at c. 22 Ma, followed by deltas that prograded towards the SSW (Rasmussen 2014). The main source area for these lower Miocene sand deposits was within southern Norway and southwestern Sweden (Olivarius et al. 2014). Olivarius et al. (2014) also found that the high maturity of the heavy-mineral assemblages of the Miocene sand, a relatively large zircon content and paucity of feldspars, indicates weathering and kaolinization in the provenance area. It could also indicate that many of the minerals were from redeposited sediments, a suggestion supported by the presence of Archean-age zircons whose nearest source area is in Finland (Olivarius et al. 2014). We suggest that the sources for the lower Miocene sands that reached the Danish Basin from Norway consisted of both sedimentary covers as well as the deeply weathered bedrock of the Mesozoic peneplain that was re-exposed after early Miocene exhumation.

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Fig. 10.

Outline of southern Scandinavia with evidence for two phases of Neogene uplift beginning between 23 and 21 Ma and at c. 4 Ma. Neogene uplift began in the early Miocene within and north of the Sorgenfrei–Tornquist Zone. However, Neogene uplift began in the Pliocene south of the zone, where Miocene sediments had accumulated in the Danish Basin. Sources: early Miocene uplift of southernmost Norway, this study; offshore SW Norway, Stoker et al. (2005) and Japsen et al. (2010); southern Sweden, Japsen et al. (2016); the Sorgenfrei–Tornquist Zone, Japsen et al. (2007, 2016); early Pliocene uplift of southernmost Norway, this study; offshore SW Norway, Stoker et al. (2005) and Japsen et al. (2010); southern Sweden, Japsen et al. (2016); Danish Basin, Japsen et al. (2007).

Neogene uplift began in the early Miocene within and north of the Sorgenfrei–Tornquist Zone (Fig. 10). However, Neogene uplift began in the Pliocene south of the zone, where Miocene sediments had accumulated in the Danish Basin (Japsen et al. 2007; Rasmussen et al. 2008). It thus seems likely that the Sorgenfrei–Tornquist Zone acted as a hinge zone during the early Miocene between the uplifting Baltic Shield and the subsiding Danish Basin.

Episodic burial and exhumation

The Norwegian margin is in many aspects a typical example of an elevated passive continental margin (EPCM) with a Mesozoic rift system parallel to the coast, a transition from continental to oceanic crust further offshore and elevated plateaux cut by deeply incised valleys, separated from a coastal plain (the strandflat) by one or more escarpments (Japsen et al., 2012a). In Norway, as at other EPCMs, these plateaux formed by erosion to the base level of the adjacent sea long after break-up and were then raised to their present elevations (Fig. 9). Furthermore, sediments deposited horizontally in the post-rift section now dip seaward and are truncated by post-rift erosional unconformities, and post-rift sedimentary systems prograde away from the margin (Figs 1 and 2). The many similarities between the Norwegian EPCM and EPCMs elsewhere in the world (e.g. in Brazil, southern Africa and SE Australia; Lidmar-Bergström et al. 2000; Japsen et al. 2012b; Green et al. 2013, 2018), lead us to suggest that the formation of the Norwegian margin is controlled by processes common to all EPCMs, and that these processes are probably plate-scale.

The processes responsible for the Neogene rise of the Southern Scandes are, however, not understood. That the controlling processes are tectonic rather than climatic is demonstrated by the separation between early and late Neogene onset of exhumation in southern Scandinavia by the Sorgenfrei–Tornquist Zone (Fig. 10), and by the fact that the major Neogene unconformities along the NW European margin predate the onset of widespread glaciations (Stoker et al. 2005). Rasmussen (2014) highlighted the tectonic origin of the changes that led to the formation of the early Miocene, braided fluvial system from the Southern Scandes into the Danish Basin. This system was capable of transporting clasts with a diameter of up to 4 cm several hundred kilometres into the basin, indicating a high-energy fluvial system and hence uplift of the source area.

Vertical motion forming EPCMs in general and the Southern Scandes in particular may be due either to sub-lithospheric processes (e.g. dynamic support from the asthenosphere; Pekeris 1935; Griffiths & Campbell 1990; Lovell 2010; Colli et al. 2014), or to compressive stresses that build up during changes in plate motion (Cobbold et al. 2007; Cloetingh & Burov 2010). Such motion may in turn be amplified by a flexural, isostatic response to erosional unloading of the landmass and depositional loading of adjacent basins (Molnar & England 1990), but the initial uplift of peneplains formed by erosion to the base level of the adjacent sea requires a tectonic trigger.

We have shown that the southernmost Scandes formed as a result of three Mesozoic–Cenozoic episodes of burial and exhumation. The rocks that are now at the surface have thus been at very different elevations above and below sea-level during that time span, which is in agreement with the quantitative estimates of the palaeotopography for Norway south of 62°N by Sømme et al. (2013a,b). This complex development is reflected in the overall configuration of the Southern Scandes where subhorizontal Palaeic surfaces (including Hardangervidda) cut off a tilted, Mesozoic peneplain (Figs 2 and 3). Here we have shown that the main Palaeic surfaces in the southernmost Scandes formed as the result of Miocene denudation, and that a cover of Jurassic–Oligocene sediments on the flanks caused the early Miocene palaeotemperatures defined by AFTA and VR. It is thus clear that the Scandes are not the erosional remnants of the Caledonian Mountains (Nielsen et al. 2009; Pedersen et al. 2016, 2018; see also Lidmar-Bergström & Bonow, 2009; Chalmers et al., 2010; Gabrielsen et al. 2010).

Erosion and isostatic adjustment of the Caledonian Mountains

Nielsen et al. (2009) hypothesized that the topography of the Southern Scandes represents the combined effects of continued erosion since the Caledonian Orogeny as a result of climatic deterioration, magnified by isostasy supported by a crustal root under the Scandes. This argument was continued by Pedersen et al. (2016), who argued that the spatial offset between the small root observed by Stratford & Thybo (2011) and the highest topography could be accounted for by flexural rigidity of the crust. Chalmers et al. (2010) showed, however, that a root 2–3 times the size of that observed by Stratford & Thybo (2011) would be necessary to support the southern Scandes. In addition, no root whatsoever was observed under the Northern Scandes by England & Ebbing (2012), although a transition to a high-velocity lower crust was observed east of the Scandes. It is true that the low-lying topography to the east, under Sweden and Finland, is underlain by thicker crust and the deep crust probably contains a high proportion of high-density eclogite (Kozlovskaya et al. 2004). Eclogite is also known from exhumed lower crust near Sognefjord in west Norway (Milnes et al. 1997), so there is likely to be eclogite in unknown proportions under all the Scandinavian crust, and, as the distribution of eclogite is unknown, arguments based on isostasy are inconclusive. Pedersen et al. (2016) did not consider that dynamic topography is reflected in the gravity signal in a complicated, wavelength-dependent manner (Colli et al. 2016), and they consequently arrived at models that are at odds with observations available in the literature and with those presented in this study.

The ideas presented by Nielsen et al. (2009) have also been criticized on a number of other grounds including geomorphological, geophysical, thermochronological, sedimentological and basic geological arguments (Lidmar-Bergström & Bonow 2009; Chalmers et al. 2010; Gabrielsen et al. 2010). We therefore do not think that arguments based on isostasy can be used to come to any definitive conclusion about the presence of the Norwegian mountains, nor any other mountains near so-called ‘passive’ margins. We need to look for other evidence, and we consider that this paper is a contribution to such evidence.

More recently, Pedersen et al. (2018) used inverse landscape evolution modelling in an attempt to constrain the topographic evolution of Scandinavia between 54 and 4 Ma, based on sediment volumes west of southern Norway and reconstructed pre-glacial topography. Their preferred model involved the presence of high (2 km) topography since 54 Ma. However, those researchers failed to consider several simple observations that contradict their conclusions and render their modelling exercise meaningless, including the following. (1) The maturity of Jurassic sediments reflects that they have been buried by a significant thickness of cover (Fossen et al. 1997), as supported by the erosional truncation of Jurassic and younger sequences towards the coast of Norway. (2) Major Paleogene and Miocene sediment accumulations occur in the Danish Basin are not included in their calculations, and notably uplift and erosion in the Scandinavian hinterland resulted in progradation of major delta systems into the basin in the earliest Miocene (Rasmussen 2014, 2017). (3) The thick Eocene sediments included in their volume calculations are primarily derived from the Shetland Platform, not from the Scandes (Ahmadi et al. 2003; Knox et al. 2010). (4) Significant volumes of sediment were removed by Neogene erosion west and south of Norway (Jensen & Schmidt 1992; Hansen 1996; Japsen et al. 2007, 2010), and unknown volumes of sediment may have been deposited east of Norway and subsequently removed. (5) The Palaeic Relief of the Southern Scandes consists of Cenozoic erosion surfaces in four steps with a vertical separation of 1500 m (Lidmar-Bergström et al. 2000, 2013), and thus the traditional interpretation of the topography of Norway as ‘remnants of a peneplain eroded to sea-level in the Mesozoic’ is incorrect (Pedersen et al. 2018).

Erosion and isostatic adjustment of Mesozoic rift shoulders

It was argued (Osmundsen & Redfield 2011; Redfield & Osmundsen 2013) that the post-rift evolution of the Scandes was driven by flexed isostatic adjustments to rift shoulders formed during Late Jurassic and Early Cretaceous rifting. They argued that accommodation of sediments above the rifted basins offshore was driven by thermal cooling, and that continued erosion of the rift-margin uplifts resulted in isostatic flexural arching of the footwall area. The uplift was sufficient that the rift-margin faults moved repeatedly, as they continue to do. Redfield & Osmundsen (2013) argued that this mechanism has acted several times in distinct episodes since the Early Cretaceous and that the stepped topography identified by Lidmar-Bergström et al. (2000) is the result of each distinct episode, but they did not, however, provide a mechanism for the episodic nature of the uplifts.

The interpretation by Redfield & Osmundsen (2013) may explain some of the effects affecting the area within and on the hanging wall of the Møre–Trøndelag Fault Complex. It is inadequate, however, to explain the uplift of the whole of the Southern Scandes. As Chalmers et al. (2010) pointed out, the width of footwall uplift of a crust with an elastic thickness of about 30 km is about 70 km, not the 300 km or more width of the Southern Scandes. Medvedev & Hartz (2014)  also showed that the flexed isostatic response of erosion of the deep valleys is of the order of 800 m, which is inadequate to explain all the modern topography.

Our AFTA data also show that post-rift subsidence and burial is not restricted to the present-day offshore domain but extended into the region that is onshore Norway today. This is demonstrated by high, early Miocene palaeotemperatures of rocks along the coastal zone where Upper Jurassic sediments rest on basement (Figs 1a and 2). The present-day, seaward facing escarpment along the west coast of the Southern Scandes is thus not a modified, Mesozoic rift shoulder, as suggested by Redfield & Osmundsen (2013). It is, instead, a Neogene feature and the remnant landscape on it (hilly relief) was formed prior to burial below late Mesozoic sediments and subsequently exhumed. This conclusion is supported by the quantitative estimates of Sømme et al. (2013b), who found that the land surface along the western margin of southern Norway did not exceed 200 m a.s.l. during the early Eocene. In fact, much of southern Scandinavia may have been below sea-level in the Eocene (Knox et al. 2010).

Dynamic topography

Rickers et al. (2013) used full waveform seismic tomography to map a low-velocity layer extending from the Iceland and Jan Mayen hotspots to beneath the continental lithosphere of the Southern Scandes, the Danish Basin, part of the British Isles and eastern Greenland, and they noted that these regions experienced uplift in Neogene times. Schoonman et al. (2017)  argued that this spatial correlation between the low-velocity layer and uplifted regions indicates dynamic support by low-density asthenosphere. However, they did not comment on the apparent absence of low-density asthenosphere under, for example, West Greenland and the Barents Sea where Neogene uplift is well documented (Bonow et al. 2006; Japsen et al. 2006; Green & Duddy 2010). Japsen et al. (2014) argued that the very high elevations in the parts of East Greenland nearest Iceland (in contrast to the margin further north and south as well as West Greenland) were the result of Pliocene uplift owing to dynamic support from the adjacent Iceland Plume superimposed on whatever process was causing the other EPCMs around Greenland. Thus, although we find it possible that dynamic support by low-density asthenosphere originating from the Iceland Plume may have contributed to the Pliocene uplift of the Southern Scandes, we also want to point out that the presence of similar EPCMs in, for example, eastern Australia and eastern Brazil cannot be explained by influence from nearby plumes.

The magnitude of the late Cenozoic (Pliocene) uplift phase in southern Scandinavia declines rapidly away from the Atlantic margin (Lidmar-Bergström et al. 2013): from 800–1200 m in the Southern Scandes to 150 m for the South Swedish Dome. This is in contrast to the similar amounts of section removed during early Miocene exhumation across southern Scandinavia, as illustrated by similar early Miocene palaeotemperatures from southernmost Norway to southern Sweden in samples at elevations close to sea-level (Japsen et al. 2016). In contrast to Scandinavia, Neogene uplift of the margins of East and West Greenland did not begin in the early Miocene, but in the late Miocene (c. 10 Ma; Japsen et al. 2006, 2014). Because the processes causing these vertical motions are not understood, it is not clear if the differences in magnitude and timing of major episodes in different areas are due to different processes or are controlled by varying responses of the lithosphere in different areas. Further insights into these questions must await improved insights of the underlying geodynamic processes.