Structure of the Scandes lithosphere from surface to depth

Tectonophysics 536–537 (2012) 1–24

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Review Article

Structure of the Scandes lithosphere from surface to depth J. Ebbing a, b,⁎, R.W. England c, T. Korja d, T. Lauritsen a, O. Olesen a, W. Stratford e, 1, C. Weidle f, 2 a

Geological Survey of Norway, Trondheim, Norway Department of Petroleum Engineering and Applied Geophysics, NTNU, Trondheim, Norway c Department of Geology, University of Leicester, United Kingdom d Department of Physical Sciences, University of Oulu, Finland e Department of Geography and Geology, University of Copenhagen, Denmark f Department of Geosciences, University of Oslo, Norway b

a r t i c l e

i n f o

Article history: Received 22 November 2010 Received in revised form 31 January 2012 Accepted 13 February 2012 Available online 25 February 2012 Keywords: Scandes Potential fields Petrophysics Seismic Lithosphere Isostasy

a b s t r a c t In this review we give an overview of geophysical data and models available for the Scandinavian mountains and adjacent areas, as they are of relevance to the debate about the existence and cause of Neogene uplift. Emphasis is given to potential field and petrophysical data of which the earth science communities of Norway and Sweden have a long tradition of acquisition and interpretation. This is reflected in the wealth and dense coverage of data available. The topography of the Scandes mountain chain correlates to a large degree with a Bouguer gravity low, which suggests isostatic compensation. But comparison with magnetic and petrophysical data shows that the gravity low is partly influenced by the Trans-Scandinavian Igneous Belt. Signals due to the effects of surface geology can be further eliminated by considering the near-surface density distribution as provided by petrophysical sampling. To illuminate the entire lithospheric structure, active and passive seismic and magnetotelluric data can be used. Recent and ongoing experiments are aimed at improving the existing models of the lithosphere. Integration of these different geophysical data sets allows the structure of the lithosphere of the Scandes and its surroundings to be addressed. Seismic models show no pronounced crustal root below the Scandes which could provide isostatic compensation, but instead an increase of crustal thickness towards the central Fennoscandian Shield. From the integrated model, isostatic compensation is inferred to be largely controlled by the density distribution of the crust and that a high-density lower crust to the east of the Scandes is necessary to achieve isostatic equilibrium on a regional scale. We demonstrate the typical crustal structure with profiles crossing from the Scandes into the shield, and a combined onshore–offshore basement map. © 2012 Elsevier B.V. All rights reserved.

Contents 1. 2. 3.

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Geological context: From the Caledonian orogeny to the uplift of the Scandes 2.1. The controversy regarding Neogene uplift . . . . . . . . . . . . . Geophysical data sets . . . . . . . . . . . . . . . . . . . . . . . . . . 3.1. Magnetic data . . . . . . . . . . . . . . . . . . . . . . . . . . 3.2. Gravity data . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.2.1. Free-air anomaly . . . . . . . . . . . . . . . . . . . . . 3.2.2. Bouguer anomaly . . . . . . . . . . . . . . . . . . . . 3.2.3. Isostatic anomaly . . . . . . . . . . . . . . . . . . . . 3.3. Petrophysics . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.3.1. Gravity effect of near-surface density distribution . . . . . 3.4. Heat-flow . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.5. Seismic studies . . . . . . . . . . . . . . . . . . . . . . . . . . 3.6. Seismological models of the lithospheric mantle . . . . . . . . . . 3.7. Magnetotelluric studies . . . . . . . . . . . . . . . . . . . . . .

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⁎ Corresponding author at: Geological Survey of Norway, Postboks 6315 Sluppen, 7491 Trondheim, Norway. Tel.: + 47 73 90 44 51; fax: + 47 73 92 16 20. E-mail address: [email protected] (J. Ebbing). 1 Now at: Department of Earth Sciences, Durham University, United Kingdom. 2 Now at: Institute of Geosciences, Christian-Albrechts University Kiel, Germany. 0040-1951/$ – see front matter © 2012 Elsevier B.V. All rights reserved. doi:10.1016/j.tecto.2012.02.016

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Lithosphere structure of the Scandes from surface to depth 4.1. Top basement estimates on- and offshore . . . . . 4.2. Crustal structure of the Scandes . . . . . . . . . 4.2.1. Southern Scandes . . . . . . . . . . . . 4.2.2. Central Scandes . . . . . . . . . . . . 4.2.3. Northern Scandes . . . . . . . . . . . . 4.3. An isostatic model for the Scandes . . . . . . . . 4.3.1. Local and regional isostasy . . . . . . . 4.3.2. A 3D isostatic model of the Scandes . . . 5. Conclusions and outlook . . . . . . . . . . . . . . . . Acknowledgements . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . .

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1. Introduction The Scandes mountain range is located at the western edge of the Fennoscandian Shield and extends in a north–south direction for over 1400 km (Fig. 1). The Scandes are the remnants of the Caledonian orogeny and its post-orogenic collapse that formed the paleo-shape of the mountains (e.g., Andersen, 1998). Tectonically the Scandinavian landmass has an intracratonic position and is in close proximity to the passive continental margin along the northeast Atlantic. As part of the Fennoscandian Shield, the Scandes were also affected by its tectonic evolution. Ongoing research is aimed at understanding, why we still see today, a long time after the demise of the Caledonides, such prominent relief, and what role more recent (e.g. Neogene or Tertiary) uplift processes have played in shaping topography (e.g., Redfield et al., 2005b; Rohrmann et al., 2002). Often, the Scandes are subdivided by their topographic shape as they feature two dome-like areas of high mountains and plateaus: the Northern and Southern Scandes (Fig. 1). While the Northern Scandes have an east–west extension of 200 km, the Southern Scandes have a maximum east–west extension of up to 400 km. In between these two areas the Scandes mountain range is narrow and not as pronounced.

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The geological structure and tectonic evolution of the western Fennoscandian Shield are also expressed in the range of geophysical data interpretations and models available. In this paper, we aim to provide an overview of the geophysical data sets and models available. Since the early 1980s most geophysical research in Norway was, due to economic reasons, concentrated on the Norwegian shelf, while in Sweden projects like the FENNOLORA profile as part of the European GeoTraverse (e.g. Guggisberg et al., 1991) were undertaken to elucidate the structure of the central Fennoscandian Shield. The ongoing debate about the processes that shaped the Scandes has stimulated a series of new geophysical experiments in recent years (e.g. England and Ebbing, in press; Stratford et al., 2009; Svenningsen et al., 2007; Weidle et al., 2010), which aim to close the gap between the wellstudied Norwegian shelf and Sweden. In this connexion, we want to provide an overview of previous, recent and ongoing geophysical studies over the Scandinavian mountain chain, and discuss which elements of the crustal and lithospheric structure are expressed in the different geophysical data sets. We do not attempt to solve the ongoing debate regarding the existence and cause of Neogene uplift (Chalmers et al., 2010; Lidmar-Bergström and Bonow, 2009; Nielsen et al., 2009, 2010), but provide a basis for evaluation of such concepts. 2. Geological context: From the Caledonian orogeny to the uplift of the Scandes

Fig. 1. Topography of the Scandes and surrounding region. The white dotted lines depict the Northern and Southern Scandes. The 500 m and 100 m contours are outlined. The yellow lines indicate the ongoing vertical uplift (in mm/year; Vestøl, 2006).

The present shape and height of the Scandes is a product of multiple tectonic events covering a timeline from the formation of the Fennoscandian Shield as part of the East European Craton in the Precambrian through the opening of the North Atlantic Ocean to recent post-glacial processes. In the following, we provide a brief summary of the main geological events. More details on the complex tectonic history and geology can be found in Gee and Sturt (1985), Ramberg et al. (2008), Roberts (2003) and references therein. The high topography of the Norwegian mountains is the remnant of a series of geological events, starting with the building of the Caledonian orogen, overprinted by late-Caledonian post-orogenic collapse, two rifting phases and possibly a period of localised Tertiary uplift. Following the Caledonian orogeny, post-orogenic collapse formed the paleo-shape of the Scandes and the passive margin system (e.g., Andersen, 1998). Onshore mid-Norway a series of dominantly extensional detachment zones can be observed which are related to postorogenic processes and may have controlled mass transfer from different segments of the Scandinavian mountain chain to the margin during the post-orogenic collapse phase (Braathen et al., 2002; Olesen et al., 2002). The pre-Caledonian basement dominates the surface geology of the Scandes (Fig. 2) and is divided by Caledonian age thrusts, normal faults and detachments. Most crustal units in the study area were created in several phases between 1.9 and 1.5 Ga (Gorbatschev and Bogdanova, 1993; Lahtinen et al., 2008). In the Western Gneiss

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Fig. 2. Simplified onshore–offshore geological map of the Scandinavian North Atlantic passive margin (modified from Mosar, 2003). The onshore tectonostratigraphic map is a simplified and modified version of the map by Gee et al. (1985). HD: Høybakken Detachment, KD: Kollstraumen Detachment, NSZ: Nesna Shear Zone, MANUS: Mandal–Ustaoset Fault, RIP: Rogaland Igneous Province, PKF: Porsgrunn–Kristiansand Fault.

Complex in southwestern Norway, the Precambrian basement is exposed as a result of transfer of Caledonian basement towards the margin during extensional collapse and erosion of the Caledonian nappes. At the eastern margin of the Svecofennian domain, the TransScandinavian Igneous Belt (TIB) is located. The TIB is a more than 1200 km long magmatic belt extending from SW Sweden to NE Norway (e.g. Olesen et al., 2010a). The TIB consists of large massifs of granitoid rocks and associated inter mediate and mafic rocks. The TIB formed during a phase of magmatic activity from 1.83 to 1.65 Ga in the westernmost part of the Svecofennian domain (Gorbatschev and Bogdanova, 1993; Högdahl et al., 2004). The Sveconorwegian orogeny at 1.15–0.9 Ga caused reworking of crustal units to the southwest. Later, a triple junction developed west of the study area as the microcontinent Avalonia was docking with Baltica around 440 Ma, followed shortly after by the collision of Baltica and Avalonia with Laurentia around 420–400 Ma. This resulted in the eastwards overthrusting of basement, overlying Baltic continental margin and

oceanic sediments over autochthonous basement in a thick series of nappes or allochthons, creating the Caledonides (Roberts, 2003). In the Carboniferous–Triassic the Oslo Graben, another important element of the geology of Scandinavia, was formed by rifting and intrusion of igneous rocks. Two major fault trends, NNW–SSE (to N–S) and NE–SW, define the borders of the Oslo Rift, which narrows towards the north. Permian age dykes, striking parallel to the major structures, are mapped far outside the graben region and indicate a strong dependency on the location of Permian intrusions upon older zones of weakness, such as Precambrian faults and fracture zones. Different models for the internal structure of the crust below the Oslo Rift exist, but a high amount of magmatic material is required to satisfy most observations (e.g., Ebbing et al., 2007; Olsen et al., 1987; Ramberg, 1976; Wessel and Husebye, 1987). Glaciations are the most recent process that carved the impressive fjord-landscape that Norway is well known for (e.g., Smelror et al., 2007). Glacial processes have been proposed as a contributing factor

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in shaping the overall topography (Nielsen et al., 2009), and deglaciation of the northern hemisphere results in the prominent postglacial rebound observed today (e.g. Lambeck et al., 1998).

2.1. The controversy regarding Neogene uplift In the present, Fennoscandia and adjacent areas are largely affected by high uplift rates (Fig. 1). The rates of present vertical uplift in Fennoscandia range from close to zero along the Norwegian coast to more than 8 mm/yr in central parts of the Gulf of Bothnia. Over some areas of the Scandes a significant vertical component of 1–4 mm/yr is observed (e.g., Milne et al., 2001, 2004; Vestøl, 2006). The main cause of the present uplift is generally assumed to be post-glacial rebound (e.g., Balling, 1980; Lambeck et al., 1998; Lidberg et al., 2010; Marquart, 1989; Milne et al., 2001, 2004; Niskanen, 1939). However, Fjeldskaar et al. (2000) argued that they were able to identify a present-day tectonic uplift component within the post-glacial rebound pattern, which coincides with thermochronologically defined areas of Neogene uplift (Redfield et al., 2005a,b). Milne et al. (2001, 2004) found no evidence for ongoing tectonic uplift on a regional scale and ruled out any ongoing horizontal tectonic motions greater than 1 mm/year. Nevertheless, in between the post-orogenic collapse and the ongoing post-glacial rebound, the Scandes may have been partly uplifted during Plio-Pleistocene time (Faleide et al., 2002; Riis, 1996; Rohrmann et al., 2002), and it has been proposed that the Scandes are part of a circum Atlantic belt of Neogene uplifts including mountains in Scotland, Svalbard and East Greenland. (e.g. Japsen and Chalmers, 2000). Rohrman and van der Beek (1996), Riis (1996) and Lidmar-Bergström et al. (2000) proposed a Neogene uplift of more than 1000 m in southern Norway based on apatite fission track data, extrapolation of the offshore late Tertiary stratigraphy and modelling of geomorphology. Similarly, Riis (1996) and Hendriks and Andriessen (2002) argued for a Neogene bedrock uplift of more than 1000 m in the Lofoten–Vesterålen area, and 600 m on the mainland to the east. For the mechanism of this Neogene uplift a variety of processes have been proposed, e.g. asthenospheric diapirism (Rohrman and van der Beek, 1996), marginal flexure and associated faulting (Osmundsen et al., 2010; Redfield et al., 2005b), but none is as yet generally accepted (see Gabrielsen et al., 2005 for a more complete review of possible processes). Nielsen et al. (2009) on the other hand argue that the present-day topography, at least for the Southern Scandes, may be equally explained by a model of protracted exhumation of topography since the Caledonian orogeny. Exhumation occurred by gravitational collapse, continental rifting and erosion. Initially, tectonic exhumation dominated, although erosion rates were high (Nielsen et al., 2009). The subsequent demise of onshore tectonic activity allowed slow erosion to become the dominating exhumation agent. The elevation limiting and landscape shaping of glacial and periglacial processes gained importance at the Eocene–Oligocene boundary when erosion rates increased. The longevity of western Scandinavian topography is explained by the failure of rifting processes to destroying the topography entirely, and by the buoyant upward feeding of replacement crustal material commensurate with exhumation unloading (Nielsen et al., 2009). In this hypothesis an additional Neogene uplift phase is not required to explain the present-day topography. This model is disputed and sparked an ongoing discussion about the validity of existing data and models (Chalmers et al., 2010; Lidmar-Bergström and Bonow, 2009; Nielsen et al., 2009, 2010). As for the models attempting to explain a Neogene uplift phase, the model by Nielsen et al. (2009) does not address the differences in topography between north and south along the Scandinavian mountain chain, which might hold a key to the understanding of its evolution.

Additionally, in a recent study, Osmundsen and Redfield (2011) compared passive margins on a global scale, and state that the topography of passive margins does not primarily reflect age, magmatism, climate or mantle convection, but more probably the response to loading of the laterally variable, fault-controlled architecture of the crystalline crust. Without attempting to resolve the ongoing debate, the following chapters present the available data and models for the present-day lithospheric structure. Analysis of the gravity field and seismic data provide a means of studying structural differences within the Scandes. A better image of the lithosphere below the Scandes and adjacent regions will substantiate the ongoing debate and help to evaluate the proposed mechanism of exhumation of the mountain range and of distinguishing between different phases of mountain shaping. 3. Geophysical data sets Over the Scandes a series of geophysical experiments and studies have been conducted, especially since the mid 1960s. Today, wellestablished databases exist for gravity and magnetic measurements for Sweden, Norway and its continental shelf. The closely spaced gravity and magnetic data points are a valuable resource for studying the deep structure of the Scandes when used in combination with petrophysical and seismic data. In Norway, the beginning of hydrocarbon exploration led to a noticeable shift from geosciences on land to commercial and scientific studies offshore. Consequently large data sets were collected offshore during the 1980s and 1990s. In the following we will summarise the regional geophysical data available for the Scandes and its surrounding areas. Hereby, we first give an overview of the data sets available for the different geophysical methods, before we try to integrate them in an interpretation of the lithospheric structure of the Scandes. 3.1. Magnetic data Systematic measurements of the magnetic field over Norway and Sweden started in the late 1950s. The Geological Survey of Sweden (SGU) started in the early 1960s with aeromagnetic measurements (e.g. Werner, 1963) and today c. 90% of mainland Sweden is covered with aeromagnetic measurements (www.sgu.se). Large areas of Norway and its surroundings were covered between 1959 and 1968 with low flight altitude (150–300 m) aeromagnetic data. From 1968 to 1976 western Norway and most of northern Norway were covered with measurements made at flight altitudes ranging from 800 m to 1500 m a.s.l. (Olesen et al., 2010a). In 1970 another high-altitude aeromagnetic survey was flown as part of the Norwegian Geotraverse Project (Aalstad et al., 1977; Heier, 1977), with a constant flight altitude of 3400 m a.s.l. and a line spacing of 3 km to reduce the effect of varying terrain and to map regional-scale features. For a complete review of aeromagnetic surveys over Norway and its continental shelf see Olesen et al. (2010a). Today, low-altitude aeromagnetic surveys are flown and measured in Norway and Sweden with a typical line spacing of 200–800 m and a flight altitude of 30–60 m. Most of these data have been included in a compilation by Korhonen et al. (2002b). Mainland Sweden and offshore Norway are covered today with higher resolution data than mainland Norway, and recent programmes at the Geological Survey of Norway aim to improve the data quality for mainland Norway with special interest for mineral prospecting. Olesen et al. (2010a,c) presented a new compilation of all available data for the Northeast Atlantic and the western part of the Fennoscandian Shield (Fig. 3). The combined magnetic anomaly map is not corrected for the different flight heights of the different surveys, and over the Scandes is partly based on data with a generally wide line-spacing, which prohibits very detailed analysis. Helicopter

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Geological Surveys of Sweden and Norway, and foreign, Swedish and Norwegian institutions. Most of these data have been collected since the end of the 1960s, using different means of transportation (e.g., cars, helicopters and boats) to reach even the most remote areas. The gravity maps presented in Fig. 4 are based on a compilation of Fennoscandian gravity data by Korhonen et al. (2002a) with adjustments for the Norwegian Sea and mainland Norway by Olesen et al. (2010a,b). For the interpretation of the gravity field, free-air, Bouguer and isostatic anomalies are commonly used (Figs. 4 and 5) as they can enhance different features. 3.2.1. Free-air anomaly Free-air anomalies show often a strong topographic signal in addition to the crustal signal. Over the Scandes a free-air gravity high (>90 mGal) is visible over the Southern Scandes, while for the Central Scandes the free-air gravity signal is small and in the Northern Scandes a gravity minima (b −50 mGal) appears just west of the highest topography (Fig. 4A). This might already indicate different segments of the Scandes mountain chain, and that the topographic signal is compensated in the Central and Northern Scandes by a negative gravity signal from within the crust or lithosphere, while in the Southern Scandes the topographic signal is dominating.

Fig. 3. The magnetic anomaly map of Norway and surrounding region is based on the compilation by Olesen et al. (2010a,c). TIB: Trans-Scandinavian Igneous Belt; OG: Oslo Graben. Black dotted lines indicate location of profiles A, B, C as shown in Figs. 13–15.

surveys with a nominal higher resolution are available for some local areas, but do not provide coverage for regional interpretations. The combined dataset from the mainland and the shelf area shows that the bedrock structures are continuous from the Baltic Shield under the Scandes and into the continental shelf (e.g., Olesen et al., 2002; Skilbrei et al., 2002). These data provide important constraints on the interpretation of the regional basement configuration and distribution of volcanic rocks on the shelf. The interpretation of aeromagnetic data also suggests a correlation of onshore detachment zones with the margin geometry offshore Mid-Norway (e.g., Olesen et al., 2002; Skilbrei et al., 2002). In the magnetic map (Fig. 3) outstanding features are, for example, the prominent high along the Lofoten islands and in the Oslo Graben. In the Oslo Graben the magnetic anomaly reaches values up to 1500 nT which can be related directly to the Permian magmatic rocks (e.g., Ebbing et al., 2007). The Trans-Scandinavian Igneous Belt can be traced as a magnetic high on the anomaly maps. This high with amplitudes in general >300 nT extends on the magnetic anomaly map from SW Sweden to NE Norway. The TIB is exposed in the south and is interpreted to continue beneath the Caledonian nappes where it intersects the Scandes (e.g. Dyrelius, 1980; Henkel and Eriksson, 1987; Olesen et al., 2002; Skilbrei et al., 2002). 3.2. Gravity data The gravity data set for the Scandes and adjacent areas is a compilation of terrestrial gravity points for Sweden and Norway, and ship-borne gravity surveys on the Norwegian shelf. For Sweden and mainland Norway respectively, more than 180,000 and 68,000 gravity measurement points exist, providing a dense sampled network over both countries. The gravity data on land have been collected by the Norwegian and Swedish Mapping Authority (Statens kartverk, Lantmäteriet), the

3.2.2. Bouguer anomaly The Bouguer anomalies are reduced to remove the effect of the topographic masses based on an onshore rock density of 2670 kg/m3 and an offshore reduction density of 2200 kg/m 3 thus enhancing the anomalies related to structures in and at the base of the crust (Fig. 4B). The lower offshore reduction density is used to minimise the density contrast between the reduced water column and the underlying low-density sediments, and to avoid an artificial signal reflecting the sea-floor bathymetry. The high topography of the Scandes mountain chain correlates with a Bouguer gravity low (in generalb −60 mGal), which is almost continuous from north to south. The general correlation between the Bouguer gravity low and the topography of the Scandes points to a form of isostatic compensation, but the more complex free-air anomaly might indicate a different degree of compensation in the different segments of the Scandes. The amplitudes of the Bouguer gravity low further suggests this as the gravity low is less pronounced in the Southern Scandes than in the Northern Scandes, which is opposite to their topographic expression. In the south the low has, in general, an amplitude of b −60 mGal, but along the Northern Scandes the gravity low becomes more circular, has a higher amplitude (−100 mGal), and is shifted to the west relative to the topography. In the south, the Bouguer low is bordered by a broad Bouguer gravity high with amplitudes up to +50 mGal close to the Oslo Graben (Fig. 4B). The irregular shape of the anomaly and its offset relative to the rift axis points to a secondary source for the anomaly in addition to the igneous rocks observed at the surface (Ebbing et al., 2005). 3.2.3. Isostatic anomaly Commonly, isostatic compensation/equilibrium means that topographic masses which sit above sea level are compensated by lowdensity material, as is the case when the crustal column below the highest topography has a lower density than the surrounding columns, if the crustal base is at a constant depth (Pratt isostasy) or a crustal root with relatively low-density material displaces highdensity mantle material (Airy–Heiskanen isostasy). For collisional orogens, Airy-isostasy is often applied; although often a combination of the two concepts explains best the geological setting, as will be demonstrated in Section 4.3.2. For a first-order interpretation, the geometry of the base of the crust can be derived from the topography if Airy-isostatic equilibrium is assumed. In addition, calculation of the isostatic residuals allows short-wavelength deviations from isostatic equilibrium to be enhanced.

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Fig. 4. A) Free-air anomaly, B) Bouguer anomaly (Reduction density: offshore 2200 kg/m3, onshore 2670 kg/m3). The gravity anomaly maps are based on the compilation by Olesen et al. (2010b). Black dotted lines in B) indicate location of profiles A, B, C as shown in Figs. 13–15.

The isostatic depth to Moho map in Fig. 5A is calculated assuming compensation by Airy–Heiskanen isostasy with a normal crustal thickness (in the absence of topography) of 30 km and a density contrast at the base of the crust of 400 kg/m 3. The resulting isostatic residual anomaly (Fig. 5B) shows, in general, anomalies with amplitudes of ±30 mGal, and often a clear correlation with the

near-surface geology. The gravity low beneath the Northern Scandes can still be observed, but with a lower amplitude (− 40 mGal). The isostatic residual shows again a clear difference between the Southern Scandes with positive anomalies up to 40 mGal and the Central and Northern part of the Scandes, where negative residuals are observed, indicating a change in the isostatic compensation along the

Fig. 5. Airy-isostasy. A) Depth to isostatic crust–mantle boundary (Airy root). The white dotted lines depict the Northern and Southern Scandes. B) Isostatic residual anomaly. The Airy root and gravity residual are calculated from the topography/bathymetry in Fig. 1 with a reference crustal thickness of 30 km and a density contrast at the Moho of 400 kg/m3. A Gaussian 150 km lowpass-filter has been applied to the Airy-root grid to smooth high-frequency variations.

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mountain belt, which probably is also expressed in the underlying crustal structure. For further insights into the residual gravity anomalies, we will, in the following section, compare the anomalies to the density distribution at the surface as estimated from petrophysical samples.

3.3. Petrophysics Some of the pronounced magnetic and gravity anomalies along the Scandes correlate with the surface geology. Consequently, it is important to know the petrophysical properties of the rocks on land when assessing the effect of near-surface structures and tectonic provinces on the potential field anomalies (e.g. Galitchanina et al., 1995). In Norway and Sweden, the Geological Surveys have carried out petrophysical sampling programmes (e.g., Elming, 1980, 1988; Heier, 1977; Henkel, 1976, 1991a; Olesen et al., 1991, 1997; Skilbrei et al., 1987, 1991) to assist the bedrock mapping and mineral exploration programmes. In addition, density measurements have been carried out on bedrock samples in conjunction with gravity campaigns (e.g., Chroston, 1974; Lønne and Sellevoll, 1975; Ramberg, 1976). Over the Norwegian mainland, approximately 28,000 rock samples collected during geological mapping and geophysical studies have been measured for density, susceptibility and remanent magnetisation. From these samples, maps of susceptibility and density have been produced (Olesen et al., 2010a; Fig. 6A, B) by calculating the average values within each geological unit represented on the Norwegian part of the northern Europe bedrock map (Sigmond, 2002). In Sweden, information about the petrophysical properties for more than 75,000 samples exist (www.sgu.se), noticeably with a gap in data coverage along the Norwegian border in Central and Northern Sweden. In combination with the compilation of the gravity and magnetic data of Fennoscandia, also a petrophysical data set has been compiled (Korhonen et al., 2002a,b). In this compilation the available samples are represented with a moving average of 15 × 15 km (Korhonen et al., 2002a). Similar to the maps for Norway, we produced from these data sets maps of density and magnetisation (Fig. 6C, D) by calculating the average values within each geological unit of the northern Europe bedrock map (Sigmond, 2002). For magnetic properties, the Fennoscandian data compilation (Fig. 6C) presents total magnetisation values, while for Norway, we present the magnetic susceptibility. Measurements of remanent magnetisation of the individual samples show large variations, and are only available for a subset of the Norwegian samples. Albeit these differences, the maps show similar features in the representation of the magnetic rock properties. Magnetisation or magnetic susceptibility can locally be high within the crystalline bedrock (Fig. 6A, C). The magmatic rocks of the Oslo Graben show higher magnetisation (susceptibility > 0.01 SI) than the surrounding bedrock, but also the Western Gneiss Region, the Precambrian basement in SE Norway and the Lofoten islands have high values. The Jotun nappes in central southern Norway, which consist mainly of mafic Precambrian rocks have a susceptibility of >0.03 (SI). Where sedimentary cover is present (e.g. sparagmites north of Oslo) the magnetic susceptibility is very low (b0.0003). The Caledonian nappes are in general magnetically transparent (compare Figs. 2, 3, and 6A, C). The density maps (Fig. 6B, D) are more comparable, and the difference between the data sets is mainly explained by the data coverage over the Norwegian mainland. In general, the two compilations have a very low mean difference of 2 kg/m 3 with a standard deviation of 83 kg/m 3. With a wealth of new data available, efforts to integrate and homogenise data from Norway and Sweden will potentially be rewarding.

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The near-surface density distribution is predominantly higher than the standard Bouguer density of 2670 kg/m3. The schists and sandstones of eastern Southern Norway have a density in the order of the Bouguer density, otherwise the Caledonian nappes which dominate the near surface geology have, in general, a density of >2750 kg/m 3. The igneous rocks of the Oslo Graben are visible as a low-density area (around 2600 kg/m 3), while the Jotun nappes in central southern Norway, which consist mainly of mafic high-density Precambrian rocks, have a high density (>2825 kg/m3) (Fig. 6D). The newer Norwegian compilation shows a higher average value >2950 kg/m3 (Fig. 6B) and the nappes east of the Jotun nappes also show a large deviation in between the two compilations, e.g. Olesen et al., 2010a: ~2700 kg/m 3; Korhonen et al., 2002a: 2830 kg/m 3. This is due to the improved sample coverage, particular for the Southern Scandes.

3.3.1. Gravity effect of near-surface density distribution Fig. 7 shows the Bouguer and isostatic anomaly reduced by a variable surface density model. Instead of the standard value of 2670 kg/m3, the density distribution (Fig. 6 B and D) has been applied from the surface to the sea-level in the calculations. The resulting gravity effect has a maximum amplitude of b30 mGal, but the largest effect in the areas of highest topography (Fig. 7C). The use of a variable surface density reduces part of the short-wavelength signal in the Bouguer and isostatic anomalies (compare Figs. 4B, 5B and 7A, B). With respect to the deeper crustal structure of the Scandes, the new gravity maps show two interesting features. The circular anomaly in the Northern Scandes is still visible in the new gravity maps, while in Southern Norway a clear transition between the Western Gneiss Region (positive residuals) and the central Caledonian nappes (negative residuals) is now visible. This demonstrates that the near surface–density distribution has a significant influence on the gravity signal. Clearly, this is a simplification as the depth extent of surface geology is more complex, and the remaining residuals can only be addressed by detailed modelling of the upper crust and the depth extent and dip of the surface bedrock units. Interestingly, the isostatic gravity residual in Fig. 7B shows a different pattern for the Southern Scandes compared to the Northern and Central Scandes, which again points to differences in the crustal structure along the Scandes mountain chain.

3.4. Heat-flow Although less data exist for heat-flow variations over Fennoscandia, interest in these data sets has increased due to economic and academic reasons, and they are here included to complete the overview. Slagstad et al. (2009 and see electronic supplement) presented a surface heat-flow compilation based on nearly 1000 onshore and offshore measurements covering Fennoscandia and the Norwegian– Greenland Sea. The data were obtained by a number of techniques, including onshore and offshore borehole measurements and probe measurements in lakes and offshore. About 60% of the borehole observations onshore Fennoscandia originate from depths of less than 500 m, and only 10% of the measurements are from depths greater than 1 km (Slagstad et al., 2009). Observations from greater depths (down to 3–5 km) are only available for the deep sedimentary basins beneath the Norwegian shelf and in the Danish areas. There is a close correlation between geological setting and heatflow. The Archean provinces show low heat-flow at ~ 35 mW/m 2. Over the Scandes heat-flow is ~50–60 mW/m 2, reflecting younger crust. The TIB can be associated with a peak in heat-flow at ~65 mW/m 2. Pascal et al. (2007) presented a model across the central Scandes, where they discussed the influence on the geometry and thickness of the TIB on surface heat-flow, and showed that heat flow data can provide additional constraints for crustal models (see more detailed discussion below).

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Fig. 6. Petrophysical maps. A) Susceptibility and B) density of surface rocks for Norway after Olesen et al. (2010c). Non-representative samples (e.g., sulphide mineralisation, hydrothermal alteration, mylonites, diabase, eclogite) were removed from the dataset (Olesen et al., 2010a). C) Total magnetisation and D) density of samples within a moving average of 15 × 15 km over Fennoscandia modified after Korhonen et al. (2002a). Total magnetisation was calculated by a scalar sum of induced and remanent magnetisation, and induced magnetisation was calculated from magnetic susceptibility by an inducing field value of 41 A/m (Korhonen et al., 2002a). All values are averaged within each geological unit in the Norwegian and Fennoscandian part of the northern Europe bedrock map (Sigmond, 2002).

3.5. Seismic studies To investigate the crustal structure of the Scandes mountain chain a number of reflection and refraction seismic profiles have been conducted, starting in the 1970s and early 1980s (e.g., Cassell et al., 1983; Hirschleber et al., 1975; Kanestrøm and Haugland, 1971; Fig. 8). Most

of these experiments do not show a significant increase in crustal thickness below the Scandes, and also had difficulties in imaging internal crustal structures. This led to doubts about the resolution of the experiments and the findings on the deep structure beneath the Scandes. In Norway for some time more focus was given to offshore seismic experiments (e.g., Barton, 1986; Christiansson et al., 2000; Lie et al.,

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Fig. 7. Use of variable near-surface density. A) Reduced Bouguer anomaly, B) reduced isostatic anomaly, and C) gravity effect of variable near-surface density. The residual anomalies have been calculated by subtracting the shown gravity effect of a variable near-surface density distribution from the Bouguer and isostatic anomaly (Figs. 5B and 6B). The gravity effect in c) has been calculated using the density distribution of Fig. 6B over the Norwegian mainland and from Fig. 6D over the remaining areas.

1990; Mjelde et al., 1992, 1997) and only a few further attempts at imaging of the onshore structure in Norway were made until recently (e.g., Hurich, 1996; Hurich and Kristoffersen, 1988; Iwasaki et al., 1994). The vast amount of multi-channel seismic reflection data, seismic refraction data and OBS lines on the Norwegian continental margin allows detailed mapping of the offshore crustal structure. Noticeably, a clear internal crustal layer can be observed on the Møre, Vøring and Lofoten margins in the form of a lower crustal body with seismic P-velocities of >7 km/s (e.g., Mjelde et al., 2005,

2009), where it is interpreted as a combination of magmatic underplating and eclogitisation or serpentinisation of the lower crust (e.g., Mjelde et al., 2009; Reynisson, 2010). Onshore Norway, one seismic profile crosses the Oslo Graben (Tryti and Sellevoll, 1977), which, however, does not clearly image the crustal structure. Experiments in the Skagerrak (Lie and Husebye, 1993, 1994; Lie et al., 1990) show significant reflectivity in large parts of the crust and also the upper mantle, demonstrating Sveconorwegian (and older) collision tectonics (Balling, 2000; Lie and

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Fig. 8. The Moho depth for Scandinavia after Grad et al. (2009) and modifications for Southern Norway after Stratford et al. (2009). The bold, black lines show locations of regional active seismic experiments over Norway and Sweden (after Kinck et al., 1993; Korsman et al., 1999) and OBS and wide-angle profiles on the mid-Norwegian margin (after Mjelde et al., 2005). The grey lines show the location of the MAGNUS-REX refraction profiles shot in 2007 (Stratford et al., 2009), orange triangles indicate station locations of the CENMOVE and CALAS project (Medhus et al., in press; Svenningsen et al., 2007), and blue triangles of the SCANLIPS experiments (England and Ebbing, in press). See references and within text for a more detailed description of the individual seismic experiments.

Husebye, 1993). Interpretations of these different data show that the Moho is elevated below the Oslo Rift and the crust is consequently thinned. In Sweden the European GeoTraverse and related experiments provide information on the central part of the Fennoscandian Shield (Guggisberg et al., 1991; Lund and Heikkinen, 1987). The results show a thick crust below the low topography of the central Fennoscandian Shield, which is not in agreement with the observed gravity signal. However, the lower crust has, in some areas, relatively high velocities (>7 km/s). Below the central Fennoscandian Shield, to the east of the Scandes, seismic studies generally show a high-velocity layer at the base of the crust (P-velocities > 7 km/s; Henkel et al., 1990; Korja et al., 1993; Korsman et al., 1999), similar to the lower crustal high velocity layer observed below the outer continental margin. Below the Bothnian Sea a thickness for this high density lower crust of up to 20 km is indicated (e.g., Korsman et al., 1999). Henkel et al. (1990) argue from combined density and seismic modelling for a high-density high-velocity lower crust beneath the Fennoscandian Shield, which they associate with eclogitisation. In a recent study, Kukkonen et al. (2008) provide evidence for eclogitisation of the lower crust beneath the subdued topography in Finland, where crustal thickness exceeds 50 km. The thickness of the high-velocity lower crust is less well

mapped in the transition to the Scandes, and below the Southern Scandes this layer is thin or even absent (e.g. Stratford et al., 2009). In the late 1980s and 1990s new efforts were made to unveil the crustal structure onshore and below the Scandes by VIBROSEIS profiles as part of the Central Caledonian Transect (CST; e.g., Hurich and Roberts, 1997; Hurich et al., 1989; Juhojuntti et al., 2001; Palm et al., 1991), in the Central Scandes, which was followed by the CABLES (Caledonian And Bothnian Lithosphere Elucidated by Seismics) wide-angle seismic experiment (Schmidt, 2000). The Central Caledonian Transect shows marked reflectivity and significant structures in the upper crust (Hurich et al., 1989; Juhojuntti et al., 2001; Palm et al., 1991), but between the upper crustal rocks and the Moho no clear reflections can be observed and on this basis a gradual increase of velocity with depth is often assumed (Juhojuntti et al., 2001; Schmidt, 2000). The CABLES profile provides the best constraint on the geometry and depth of the Moho but this is complicated by the presence of a near surface low velocity zone (Schmidt, 2000). The crustal thickness must increase rapidly to over 40 km between the coast to the start of the wide-angle profile under the Scandes and remains thick eastwards. No significant crustal thinning east of the high topography towards the central Fennoscandian Shield is observed, which again does not correlate with the gravity signal. Ottemöller and Midzi (2003) used permanently installed seismic stations to study the crustal thickness of Norway, and while they confirmed the overall geometry of the crust, the crustal thickness at some stations showed large deviations from previous estimates. From all these observations the question arose of where the isostatic compensation of the Scandes occurs (crust or upper mantle) and if the Scandes are at all in isostatic equilibrium. This stimulated a series of new passive and active seismic studies of the crust to unequivocally image the crustal structure of the Scandes mountain chain (England and Ebbing, in press; Stratford et al., 2009; Svenningsen et al., 2007). The differences and similarities of these studies are now presented and discussed here. Svenningsen et al. (2007) calculated receiver functions from teleseismic data collected with temporary seismic stations on two profiles in southern Norway to resolve crustal thickness. Using a velocity model derived from an OBS wide-angle/refraction study of the Sognefjord (Iwasaki et al., 1994), these authors imaged Moho depths from 29 km at the coast to 43 km below the highland plateau. Generally the depth to Moho is close to or exceeds 40 km beneath areas of high mean topography (~1 km), whereas in the Oslo Rift the crust locally thins to 32 km (Svenningsen et al., 2007). Svenningsen et al. (2007) also found a deepening of the Moho beneath the low topography of westernmost Sweden, comparable with the observations of Schmidt (2000). This experiment has been followed up by a set of three seismic refraction profiles across Southern Norway; the MAGNUS-REX experiment (Stratford et al., 2009). The interpretation of these data show that the Moho depths beneath the highest topography are around 38–40 km, and in general are close to the average for continental crust with elevations ~ 1 km. These Moho depths are ~ 2 km deeper than earlier estimates based on interpolation from coarsely spaced refraction profiles, but up to 3 km shallower than the receiver function estimates by Svenningsen et al. (2007). The differences in inferred Moho depth may, in part, be due to azimuthal variation in the receiver functions and biassed by the P-wave velocity model in the receiver function depth migration, which is 3% higher than has been found from the new refraction profiling (Stratford et al., 2009). For the deepest receiver functions, the piercing points which dominate the profile (Svenningsen et al., 2007: their Fig. 6) are located to the northeast, and conform well with the Moho deepening towards the northeast shown in Stratford et al. (2009). But differences in frequency content between the two data sets and the influence of the thickness of the Moho boundary may be other factors that contribute to interpretation differences (Stratford et al., 2009).

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The Moho depth variation beneath the mountains roughly correlates with changes in surface topography indicating that topography is, at least to the first order, controlled by crustal thickness variations (Stratford et al., 2009). But these authors state further that the highest mountains do not overlie the thickest crust and additional support for topography is likely. The Oslo Graben is again a deviation from this observation, and an increasing crustal thickness is observed towards central Fennoscandia. In the central Scandes along the location of the CABLES profile, an additional receiver function study was carried out as part of the SCANLIPS experiment (England and Ebbing, in press). The central profile is the first of two profiles to analyse the crustal structure in the Central and Northern Scandes. England and Ebbing (in press) confirm, in general, the results of Schmidt (2000), estimating a crustal thickness of 32 km at the Norwegian coast and an increase to 43 km beneath the Central Scandes, a depth which remains nearly constant beneath all stations eastwards until the Bothnian Sea. In addition, the experiment resolves an up to ~10 km thick high-velocity lower crust beneath Sweden that tapers out westwards below the highest topography. The high-velocity lower crust has not been resolved by the previous experiments, but the new results agree with the compilation by Korsman et al. (1999), and the study by Henkel et al. (1990). In summary, these experiments show a relatively consistent image of crustal thickness. Fig. 8 shows the crustal thickness map of Fennoscandia after Grad et al. (2009) with modifications for the Southern Scandes after Stratford et al. (2009), which is in general consistent with previous compilations (Kinck et al., 1993; Korsman et al., 1999). The Moho in general deepens from the Norwegian coast (~30 km) rapidly below the Scandes. To the north of the Oslo Rift, the crust generally continues to thicken eastwards from 30 km along the coast to more than 50 km beneath the central Fennoscandian Shield. In the Southern Scandes the Moho depth has more of an apparent than an actual root as the crustal thickness decreases towards the Oslo Graben before it thickens again towards the east. In Fig. 9 Moho depth estimates for Southern Norway from different sources are shown. Stratford et al. (2009) presented, based on the MAGNUS-Rex and previous experiments, a new crustal Moho depth map for Southern Norway. This map probably shows a more accurate Moho depth beneath the Southern Scandes than previous compilations (e.g. Grad et al., 2009; Kinck et al., 1993; Korsman et al., 1999; Tesauro et al., 2008) as an improved data base was available. Most of the differences between the compilations can be related to the data sets available for the compilations and also to different interpolation techniques. For example, overall, the compilation map of Moho depths of Stratford et al. (2009) for onshore southern Norway is similar to that of Kinck et al. (1993). The new data have added details and improved the reliability significantly. Stratford et al. (2009) estimate the uncertainty of Moho depth estimates to be in general in the order of ±2 km along the seismic profiles. Stratford et al. (2009) state that their compilation leaves the original Moho map by Kinck et al. (1993) unchanged along the lines of the prior studies and where there is overlap between the prior studies and the new data, there is agreement on Moho depth within the range of uncertainties. Offshore, the two compilations are more different as the data sets have been significantly changed in the last decade (e.g., Mjelde et al., 2009), and Stratford et al. (2009) do not attempt to provide a detailed image of the shelf. Kinck et al. (1993) and the compilations by Tesauro et al. (2008) and Grad et al. (2009) were aimed at providing a Moho depth for a larger region, and noticeably the latter two were compilations for Europe and were not specific to Southern Norway. The compilation EuCRUST-07 (Tesauro et al., 2008) makes use of different compilations and data sets, e.g. beneath the Southern Scandes the receiver functions study by Svenningsen et al. (2007) have been used, and has a resolution

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of 15′ × 15′ whenever supported by seismic data. Further differences can be explained by the way offshore data were integrated and the method of interpolation used. The compilation by Grad et al. (2009) is very different to the compilation of Kinck et al. (1993) and Stratford et al. (2009) for the Southern Scandes, which is again due to the data base used. For the Southern Scandes, the authors use the results by Ottemöller and Midzi (2003), and almost none of the available refraction profiles. This results in uncertainties, which are in general larger than ±2.5 km and in places up to ±4.5 km (Grad et al., 2009). Within these limits the data are consistent with the other studies. Over the remaining Fennoscandian Shield, where we will make use of this data set later, the compilation is more reliable as it is based on the compilation by Korsman et al. (1999) and Hjelt et al. (2006). These compilations make use of the same data base as Kinck et al. (1993) for southwestern Fennoscandia, but of additional data for the central part of the Fennoscandian Shield. 3.6. Seismological models of the lithospheric mantle The crustal scale seismic experiments leave no doubt that a discord exists between the observed gravity field and the crustal thickness, implying that the Scandes are not a typical mountain belt which is sustained by a crustal root. Hence there must be structures in the crust or upper mantle that explain this discord. Lund (1979) analysed the upper mantle structure below the Blue Road profile across the Northern Scandes, and stated that there is evidence for a shear (S-wave) low-velocity zone of a few tens of kilometres in the uppermost mantle. In an early regional P-wave travel time tomography (P-tomography) study (Husebye and Hovland, 1982; Husebye et al., 1986) general trends of reduced upper mantle velocities beneath southern Norway and the Norwegian–Danish Basin and higher upper mantle velocities beneath southern Sweden, were inferred. Bannister et al. (1991) analysed the uppermost mantle seismic velocities below Fennoscandia from Pn and Sn phases and imaged low sub-Moho velocities beneath the areas of high topography along the Scandes. As low velocities generally indicate lower densities, these experiments point towards a possible upper mantle contribution to the compensation of the topographic masses. These low Pn and Sn values correlate to some degree with areas of the proposed maximum Neogene uplift. These studies could not, however, give an estimate of the extent of the velocity anomalies in the mantle, and the link to surface geology remained unclear. A regional surface wave tomography model for Northern Europe confirmed a low-velocity zone in the mantle, which could contribute to the compensation of the high topography (Weidle and Maupin, 2008). The study imaged a pronounced negative anomaly in shear wave speed beneath southwestern Scandinavia which was linked to an extended low-velocity regime beneath the North Atlantic (Fig. 10). To improve the knowledge of the upper mantle structure beneath the high topography of southern Norway, these results were followed up by the temporary deployment of a seismological network in Southern Norway, the MAGNUS (MAntle investiGations of Norwegian Uplift Structures) experiment (Weidle et al., 2010). Fig. 10 summarises the presently available coverage of passive seismic data from temporary and permanent long period (>60 s) seismic stations. It should be noted that the Norwegian National Seismic Network is currently being modernised which will further improve the coverage of mainland Norway in the near future (Ottemöller, pers. comm.). In addition, several other temporary deployments of long period broadband seismometers were conducted in the region (CENMOVE and CALAS projects, Medhus et al., 2009 and additional instruments from the Danish Instrument Pool, Medhus et al., in press), which provide the possibility of analysis of a larger area. Medhus et al. (2009) presented an initial study of teleseismic Pwave travel time residuals (P-residuals), which has been followed up by a P-wave tomography of southwestern Scandinavia (Medhus

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Fig. 9. Depth to Moho for Southern Norway from the compilations by a) Kinck et al. (1993) and b) Tesauro et al. (2008), c) Grad et al. (2009) and d) Stratford et al. (2009) In c) as contour lines the uncertainties of the Moho depth are drawn as provided by Grad et al. (2009). In a) and d) the white lines show available deep seismic profiles. Note that the MAGNUS-REX profiles have been used to improve the map for Stratford et al. (2009).

et al., in press). The P-traveltime residuals show late arrivals in most of southern Norway west of the Oslo Graben area and in the deep Norwegian–Danish basin, and early arrivals in southwestern Sweden and southern Norway east of the Oslo Graben. These geographic patterns map in the tomography into velocity anomalies (slow beneath

southern Norway and Denmark, fast below Sweden) that point to a sharp mantle transition between Norway and Sweden which is also indicated by S-wave traveltime anomalies (Wawerzinek et al., 2011). This was postulated earlier by Pascal et al. (2002) as an explanation for the location of the Oslo Graben.

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decreasing lithospheric thickness from the centre of the Baltic Shield towards its edges. For example Calcagnile (1982) shows that the base of the lithosphere deepens from a level of 110 km below the Southern Scandes to 170 km below the Bothnian Sea (Fig. 11A). More recent studies differ in the absolute depth to this boundary (e.g., Artemieva and Thybo, 2008; Bruneton et al., 2004). The overall geometry of the base lithosphere is, however, consistent between the different studies, without revealing local patterns below the Northern and Southern Scandes. In summary, some evidence is available for a change in mantle properties between Southern Norway and Sweden. At least beneath the Southern Scandes a low-density (lithospheric) mantle may be present, which can provide additional support for the topography of the Scandes mountain chain. A better understanding of the interplay of deep and shallow compensation of the topography will require careful integrated modelling as will be attempted in Section 4.3. 3.7. Magnetotelluric studies

Fig. 10. Shear wave velocity at 115 km depth from the model by Weidle and Maupin (2008) and coverage of broadband (natural period ≥ 100 s) seismic instruments in the region. The permanent stations in Bergen, Hagfors, Kongsberg, as well as the NORSAR array are complementary to the temporary MAGNUS network and therefore presented in the same colour. The temporary networks are listed with their operational period. Note that the model is based on data available in 2005, thus using only part of the data available today, in particular in southern Norway. Refined models for southern Norway can be found in Maupin (2011), Medhus et al. (in press) or Köhler et al. (2011).

In the model by Weidle and Maupin (2008) this transition appears to be more gradual and misplaced (Fig. 10) which is attributed to the lower lateral resolution of surface waves and that the model is based on data available in 2005, thus using only part of the data available today, in particular in southern Norway. First results from surface wave analysis of the MAGNUS data substantiate the initial observations by Weidle and Maupin (2008) that shear wave velocities in the upper mantle under southern Norway are on average slow and the velocity–depth profile is more typical for continental than for cratonic mantle (Maupin, 2011; Weidle et al., 2010). 3D analysis of surface wave data indicate that there are only minor lateral variations in shear wave velocity in the uppermost mantle under southern Norway (Köhler et al., 2011) and the lateral extension of the MAGNUS network may be too small to image the transition to faster lithosphere under Sweden with surface waves. Models for the base of the lithosphere (the lithosphere–asthenosphere boundary) are available from surface wave and anisotropy studies over Fennoscandian Shield (Bruneton et al., 2004; Calcagnile, 1982; Pedersen et al., 1994; Plomerova et al., 2008). These studies image a

Petrophysical and structural information can also be defined by mapping of electrical conductivity and its possible anisotropy. Magnetotelluric (MT) measurements are useful for imaging both structure within the basement and the upper crust (Korja et al., 2008) as well as the deep lithosphere (e.g., Smirnov and Pedersen, 2009). Elsewhere in Fennoscandia, a large number of local and large scale MT surveys have been conducted during the last three decades (e.g. Korja, 2007; Korja et al., 1989, 2008; Rasmussen et al., 1987; and references therein); including several large international research projects in the Fennoscandian Shield (Hjelt et al., 2006; Lahti et al., 2005) and its margins (Brasse et al., 2006; Jones, 1983; Smirnov and Pedersen, 2009). These studies have resulted in a relatively detailed 3D model of the crustal structure (Korja et al., 2002) and in a preliminary map of lateral variations in the depth to the lithosphere–asthenosphere boundary in Fennoscandia (Korja, 2007). Estimates of the base lithosphere from the MT method (Fig. 11b according to Korja, 2007) show a similar trend to the shear-wave study by Calcagnile (1982) (Fig. 11). In comparison to the seismic lithosphere, which defines the base lithosphere as the boundary between the outer shell with higher seismic velocities than in the directly underlying asthenosphere, the electrical lithosphere is defined as the resistive outer shell overlying a highly conducting shell in the upper mantle, also called the electrical asthenosphere (Martinec and Wolf, 2005). The electric lithosphere shows a thickness of 250 km beneath the Central Fennoscandian Shield (Korja, 2007). The recent comparison of seismic and electric estimates of the depth to the lithosphere–asthenosphere boundary in Europe (Jones et al., 2010) shows that in the Precambrian East European Craton the electrical lithosphere is in average c. 70–80 km thicker than the seismically defined lithosphere. This is in contrast to the Phanerozoic Europe, to the south-west of the Transeuropean Suture Zone, where electrical lithosphere is c. 10 km thinner than the lithosphere according to receiver function data and 40 km thinner than lithosphere from seismic anisotropy. Towards the Scandes in the west and the Barents Sea at the northern margin the electric lithosphere thins to 150 km or below (Fig. 11), but deep electromagnetic (magnetotelluric, MT) data are missing from the area of the Scandinavian mountain range except for a recent MT profile in the Central Scandes (Korja et al., 2008). The profile crosses the region of the lowest topography, and focus of the research in its first phase was on the crustal structures. To overcome this gap in the data and in support of the ongoing seismological studies, a series of MT experiments have been established in the Southern and Northern Scandes within the ToSca–DMT (TopoScandia–Deep Magnetotellurics) project. These new installations will hopefully allow tracing of variations in the

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Fig. 11. Base lithosphere according to (A) seismic studies (Calcagnile, 1982) and (B) magnetotelluric studies (right; Korja, 2007). Contour lines indicate depth to the seismic (sLAB) and electrical (eLAB) base lithosphere. Black dots in eLAB map denote the sites used in the compilation.

lithospheric mantle from Sweden to Norway, and beneath the Scandes. Preliminary results of these new and ongoing MT experiments indicate a large contrast in resistivity between the Caledonian and Precambrian basement. Of particular importance are the highly conducting alum shales between the Precambrian basement and the overlying, generally resistive, Caledonian nappes as the mapping of alum shales yields a clear image on the geometry of the basement (Korja et al., 2008). Besides the deep lithospheric structure, the new results will assist in understanding the crustal structure beneath the Southern and Northern Scandes (Smirnov et al., 2011). 4. Lithosphere structure of the Scandes from surface to depth In the preceding section we presented the main geophysical data for the structure of the Scandes. In the following, we summarise the main findings and integrate them into a picture of the lithosphere beneath the Scandes. We present a basement map for the Scandes, illustrate the typical crustal structure in cross-sections through Southern, Central and Northern Norway, and discuss the isostatic state of the lithosphere. 4.1. Top basement estimates on- and offshore The Caledonian nappes dominate the surface geology, and their depth extent is of interest for understanding the link between surface geology and crustal structure. The seismic refraction experiments in the southern Scandes give surprisingly little information about their depth extent (e.g., Stratford and Thybo, 2011b; Stratford et al., 2009), but the basement configuration is visible in the gravity and magnetic anomalies and the petrophysical maps (Figs. 3–7). In general, the gravity and especially the magnetic field are very sensitive to near-surface sources, and therefore, the interpretation of potential

field data helps to model the depth extent of surface geological structures. Along the Central Scandes the combined interpretation of gravity and magnetic data (e.g., Dyrelius, 1980, 1985; Elming, 1980; Wolff, 1984) and seismic reflection profiles (Hurich, 1996; Hurich et al., 1989; Palm et al., 1991) produced images of the near surface geology in great detail, particularly with respect to the thickness of the Caledonian nappes. The seismic data indicated that the preserved Caledonide allochthon in central Scandinavia is 15 to 20 km thick, considerably greater than the 2 to 5 km predicted from surface geologic data (Hurich, 1996). However, recent seismological experiments indicate only a thickness of 10–12 km for the Caledonide allochthon (England and Ebbing, in press). MT data in general confirm these estimates and the geometry of the Caledonian nappes and the Precambrian basement (Korja et al., 2008). In Norway detailed reflection seismic data sets are mostly not available, and as explained previously the interpretation of newly acquired MT data is ongoing. Fig. 12 shows a tentative depth to basement maps for mainland Norway and the Norwegian shelf as compiled from different sources (Ebbing and Olesen, 2010; Henkel, 1991b; Olesen et al., 1990, 2002, 2010a; Skilbrei and Olesen, 2005; Skilbrei and Sindre, 1991; Skilbrei et al., 1987, 2002). Onshore in southern Norway the basement depth is taken as the base of the Caledonian nappes or rocks deformed by the Caledonian orogeny. In the Caledonian nappe region in the southeast the estimates of the basement depth are based primarily on magnetic depth estimates (Nystuen, 1981), whereas in other parts density modelling and geological mapping are the primary tools (e.g. Skilbrei and Sindre, 1991; Wolff, 1984). The uncertainty of the estimates is relatively high and new estimates based on the extended petrophysical data base should be made. In northern Norway, the basement consists mostly of Archean and Palaeoproterozoic rocks, such as metagranites, granulites, gneisses, metabasalts and amphibolites, and partly of Caledonian nappes. The

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Fig. 12. Basement estimates for Norway and the continental margin. Onshore, in southern Norway, basement depth is taken as the base of the Caledonian rocks or Caledonian nappes, and in Northern Norway, the basement consists mostly of Archean and Palaeoproterozoic rocks (Olesen et al., 2010a). Offshore, the top basement is defined as the base of the sedimentary strata (Ebbing and Olesen, 2010). For more details see text. Yellow lines indicate locations of profiles A, B, C as shown in Figs. 13–15.

basement depth is taken as the depth to the Precambrian basement and largely depends on magnetic and gravity field interpretations (Olesen et al., 2010a). Offshore, the top basement is defined as the base of the sedimentary strata. The top basement map of the Norwegian shelf was compiled by integrating individual regional studies (Ebbing and Olesen, 2010, and references therein). The character of the top basement surface differs between segments of the margin. The basement itself can be differentiated by petrophysics into two broad types: a highly magnetic part corresponding to Precambrian granulite facies rocks and a low magnetic part corresponding to Caledonian nappes and amphibolite facies Precambrian basement (Ebbing et al., 2009; Olesen et al., 1991; Skilbrei et al., 1991).

4.2. Crustal structure of the Scandes In this section, we will link the structures visible at the surface to the crustal structure of the Scandes by showing three general crosssections. Seismic studies, in combination with gravity anomalies, provide information about the deep structure below the Scandes. Magnetic anomalies place a further constraint on the extension of nearsurface geology to depth. To the north of the Oslo Graben, the Trans-Scandinavian Igneous Belt (TIB) can be observed at the surface and on magnetic anomaly maps from southern Sweden to the Lofoten Archipelago in northern Norway (Figs. 1 and 3). Possibly the TIB continues northward into the Barents Sea (Olesen et al., 2010a). The granitoid rocks of the TIB have relatively low densities (e.g., Skilbrei et al., 2002) and high

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heat production, which is also reflected in the heat-flow maps of Fennoscandia (e.g., Balling, 1995; Slagstad et al., 2009). From forward modelling of gravity anomalies in the northern and central Scandes it has been argued that these granitoids can extend to a depth between 10 and 20 km (Dyrelius, 1980; Olesen et al., 2002; Pascal et al., 2007; Skilbrei et al., 2002) which is a significant fraction of the crust. Addressing the interplay between the TIB and the axis of the highest topography in Norway is imperative for the understanding of the isostasy and dynamics of the mountain belt. In the following, we present three W–E cross-sections illustrating the differences in crustal structure from south to north. The crosssections (see locations in Fig. 12) are based on the data sets as outlined in the previous sections and explained below. The profiles begin in the west on the continental shelf and cross the Scandes in areas of the highest topography in the north and south and along the less prominent topography in the central part, before they end in the area of subdued topography in Sweden. Only the crustal structure is shown in Figs. 13–15 but for all the three profiles a base lithosphere has been included according to Calcagnile (1982) with a negative density contrast of 20 kg/m 3 between the lower lithosphere and upper asthenosphere. This density contrast is necessary to adjust a long-wavelength slope in the gravity signal (e.g. Bielik et al., 1996), which is also observed in the geoid undulations (e.g., Ebbing and Olesen, 2005) and which has the same trend and amplitude as the gravity signal associated with postglacial rebound (e.g., Mitrovica and Peltier, 1989; Sjöberg et al., 1994). The applied density contrast is equivalent to a geologically more correct modelling of a temperature dependent density gradient in the lithospheric mantle. Density distribution in the lithospheric mantle is largely temperature dependent and at the base lithosphere (defined as a temperature boundary) no density contrast with the asthenosphere must exist (e.g. Zeyen and Fernàndez, 1994). However, the gravity signal caused by temperature dependent mantle densities is equivalent to a gravity signal due to a minor density contrast placed at the base lithosphere. The uncertainties in this assumption affect the crustal sections only to a minor extent. More systematic analyses are needed to verify the thermal structure of the lithospheric mantle, but these results are very sensitive to estimates of the base lithosphere (e.g. Glaznev et al., 1996; Kolstrup, 2010). Density and magnetisation are defined by conversions of seismic velocities to densities (Table 1) and from the surface petrophysical data (Fig. 6). The velocity to density conversion was made using the linear velocity–density equations developed by Christensen and Mooney (1995) and corrected by Zoback and Mooney (2003) for crystalline crust. The simple model provides a reasonable fit between the modelled and observed gravity and magnetic data, and illustrates the changes in the crustal configuration at a regional scale. 4.2.1. Southern Scandes The first profile crosses offshore from the Viking Graben roughly in between the profile presented in Svenningsen et al. (2007) and the southern MAGNUS-REX line (Stratford and Thybo, 2011a; Stratford et al., 2009) through the Southern Scandes and the northern end of the Oslo Graben into Sweden (Fig. 13). The geometry of the Moho depth and crustal configuration is according to Christiansson et al. (2000) for the offshore part, Stratford et al. (2009) beneath the Southern Scandes and Grad et al. (2009) to the east of the Oslo Graben. Along this profile a relative deepening of the Moho can be observed from the Norwegian coast to the Southern Scandes, which almost perfectly coincides with the isostatic Moho depth. However, the Moho deepens further towards Sweden, interrupted only by the sharp transition to the thinned crust of the Oslo Rift. Variations in the upper 10 km of the crust are based on a variety of studies. The sedimentary thickness offshore and the thickness of Caledonian nappes are modelled according to the compilation presented in Fig. 12. The gravity high above the Oslo Rift is caused by a

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Fig. 13. Geophysical profile A across the Southern Scandes. The blue and white dashed line indicate the Airy-type isostatic Moho (Fig. 5A) and the isostatically derived top of the high-density lower crust (Fig. 16B), respectively. The topographic, magnetic and gravity data have been smoothed by averaging the data perpendicular to the profile orientation for a 10 km wide path. In the model calculations, the asthenosphere–lithosphere boundary (density contrast of − 20 kg/m3) is includes according to Fig. 11. 2D model calculations were performed using GMSYS-2D. Location of the profile is shown in Figs. 3, 4 and 12.

combination of crustal thinning and the presence of a mid-crustal ramp relating to Precambrian gneiss complexes at the western flank of the rift (Ebbing et al., 2005, 2007). The granite in Sweden has been introduced in the near-surface to fit a local gravity minimum and can be related to the southward extension of the TIB. Beneath the Oslo Graben a high-density lower crust is modelled which was proposed by early studies of the Oslo Graben (e.g., Ramberg, 1976). Along a cross-section the gravity field shows a steep, westward-facing gradient, partly located to the west of the rift, and a much gentler eastern gradient. The steep westwardfacing gradient is partly related to Precambrian gneiss complexes to the west of the graben (Ebbing et al., 2007), and partly reflects the thickness of the high-density lower crust (e.g. Stratford and Thybo, 2011a). The high-density lower crust continues in the model to the east of the graben and thickens beneath Sweden. Beneath the Oslo Graben, the high-density lower crust is often interpreted as a result of the emplacement of magmatic partial melt during rifting (e.g., Neumann, 1994; Olsen et al., 1987). The receiver function results of

Svenningsen et al. (2007) indicate a different lithology of the lower crust under the Oslo Graben, where its top is the main converter compared to lower crust beneath Sweden, where its base appears to be the main converter. Beneath the Southern Scandes, the lithospheric mantle is modelled with a slightly decreased density (~3350 kg/m 3) than the surrounding mantle on the shelf and beneath Sweden (~3370 kg/m 3). These densities are used to model the gravity field, and their isostatic implications will be discussed in Section 4.3. 4.2.2. Central Scandes The second profile (Fig. 14) crosses the Central Scandes along the location of the CABLES and SCANLIPS profiles. The crustal geometry offshore is according to the models presented by Ebbing et al. (2009) and Ebbing and Olesen (2010) and based on a wealth of seismic and well data in addition to gravity and magnetic data. Onshore the crustal geometry is based on the recent SCANLIPS results (England and Ebbing, in press). The upper crust is modelled according to Fig. 12 and the models presented in Skilbrei et al. (2002). Beneath

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Fig. 14. Geophysical profile B across the Central Scandes. See Fig. 13 caption for more details. Location of the profile is shown in Figs. 3, 4 and 12.

the highest topography the granitic rocks of the TIB are present beneath the Caledonian nappes which form the present surface. This structure has a relatively low density (2690 kg/m 3), high magnetic susceptibility (0.065 SI) and extends to >10 km depth. This is in agreement with earlier studies (Pascal et al., 2007). In addition to the upper crustal structure and the TIB, which are prominent in the magnetic anomalies, the variations of the Moho strongly influence the Bouguer anomaly pattern along the profile. The area of high topography and the associated Bouguer gravity low is narrower than on the southern profile and thickening of the crust is more rapid from the coast beneath the area of high topography. The crustal thickness exceeds substantially the isostatic crustal thickness beneath the Scandes and stays constant beneath Sweden where crustal thickness remains constant at ~ 40 km despite the absence of prominent topography. The lower crustal body has been imaged by the recent receiver function study (England and Ebbing, in press) with an increased seismic P-wave velocity (>7 km/s), a thickness of up to 10 km and a calculated density of 3200 kg/m 3. The part of the profile not associated with a high-density lower crust is also narrower than in the south and the high-density

lower crust tapers out below the Scandes and increases in thickness beneath Sweden. 4.2.3. Northern Scandes The northernmost profile (Fig. 15) across the Northern Scandes coincides with the location of the Blue Road traverse (Hirschleber et al., 1975), and its offshore extension Blue Norma (Theilen and Meissner, 1979; Weinrebe, 1981). The crustal geometry is modelled according to Olesen et al. (2002) and the basement thickness shown in Fig. 12. The main part of the negative gravity anomaly and the positive magnetic high is modelled with granitoid units of the TIB (Olesen et al., 2002). The TIB is located beneath and adjacent to the area of highest topography. Along this profile the TIB has a low density (c. 2650 kg/m 3), and high susceptibility (0.085 SI) and a depth extension of around 10 km. The topography is higher than on the central profile, but the crustal thickness increases again rapidly from the coast beneath the Scandes and further (to more than 40 km) beneath Sweden, as indicated by the crustal thickness compilations (Fig. 8). Also here, the crustal thickness exceeds that necessary to achieve isostatic compensation, and an additional

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Fig. 15. Geophysical profile C across the Northern Scandes. See Fig. 13 caption for more details. Location of the profile is shown in Figs. 3, 4 and 12.

high-density mass structure needed to achieve isostatic equilibrium and to explain the gravity signal. Also here, the high-density lower crust is present beneath Sweden and must taper out below the Scandes. Its westward limit correlates with that of the TIB. The TIB cannot be observed at the surface, but its extension from Sweden beneath this section is based on an interpretation of the magnetic anomaly (Fig. 3). The thickness of the high-density lower crust is based on an isostatic model and will be discussed below in more detail.

4.3. An isostatic model for the Scandes A sharp contrast in surface topography between the western part of the Fennoscandian Shield (i.e., Scandes in Norway and Sweden) and its internal regions (Precambrian Shield in Sweden and Finland) exists that must be reflected in the crustal/lithospheric structure. As mentioned before, the correlation between topography and Bouguer gravity low over the Scandes suggests, at least to some degree, isostatic compensation. For young collisional orogens, such as the Alps

Table 1 P-wave velocity and density for the crust and upper mantle beneath the Scandes. The values from (1) Stratford et al. (2009) and (2) England and Ebbing (in press) indicate velocities at the upper and lower boundary of the layers and a gradient in between. [#] High value is only observed beneath the Oslo Graben. [*] Value represents the possibly ophiolite rock associated with the upper allochthon (Andersen, 1998; Roberts and Gee, 1985).

Upper crust Middle crust Lower crust High-velocity/-density lower crust Upper mantle

Southern Scandes(1)

Central Scandes(2)

P-wave velocity [km/s]

P-wave velocity [km/s]

Density range [kg/m3]

Average density [kg/m3]

5.8–6.2 [6.5*] 6.3–6.4 [6.6#] 6.6–6.8 7.1 8.05

5.6–6.0 6.0–6.6 6.6–7.2 7.2 >8

2580–2775 [2880*] 2700–2920 2940–3170 3160–3200 3370

2700 2810 3070 3170 3370

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or the Himalayas, isostatic compensation largely occurs in the form of a crustal root as predicted by Airy–Heiskanen isostasy. Over Fennoscandia the isostatic compensation appears to be more complex. No clear crustal root is observed beneath the high topography of the Scandes, which are the remnants of the Caledonian orogeny; rather there is continuous thickening of the crust towards the east. Another structure, which influences the mode of isostatic compensation of the Scandes, is the lithosphere–asthenosphere boundary. This is especially important for Fennoscandia due to the response to post-glacial rebound. In all cross-sections a correction was made for a regional gradient, which is expressed as the geometry of the base lithosphere. Before we discuss the isostatic implications of the lithosphere configuration in detail, we will review some discussions concerning the degree of local and regional compensation for the Fennoscandian Shield.

4.3.1. Local and regional isostasy The Fennoscandian Shield is currently undergoing both vertical uplift and horizontal extension (e.g., Lidberg et al., 2010; Milne et al., 2001). The uplift rates are highest for the central Fennoscandian Shield (Fig. 1) and less pronounced for the Scandes. Ideally, complete isostatic models must consider the effect of post-glacial rebound and correct for the surface topography and gravity signal, accordingly. Post-glacial rebound occurs at the lithospheric scale and to what extent the lithosphere responds to loading is further controlled by its flexural rigidity. Flexural rigidity characterises the apparent strength of the lithosphere, which acts against the forces induced by loading. Most of the regions discussed above show clear evidence of dynamic support, which is expressed by the changes in the gravity field with time as caused by post-glacial uplift. This relation was established by comparison of repeated high precision gravity measurements and sea-level and levelling data (e.g. Ekman and Mäkinen, 1996), and is perhaps most obviously observed in the results from the GRACE satellite mission (e.g. Steffen et al., 2008). Estimates of the magnitude of the presently observed gravity field caused by the dynamic mass distribution due to post-glacial rebound have varied between 15–20 mGal (Balling, 1980) and 25–30 mGal (e.g., Mitrovica and Peltier, 1989; Sjöberg et al., 1994). This negative signal has wavelengths such that it is difficult to distinguish from a gravity signal related to the base of the lithosphere, but might as well reflect uncertainties about the crustal thickness estimates itself. Studies of the flexural rigidity for the Scandes and the Fennoscandian Shield (e.g., Ebbing and Olesen, 2005; Fjeldskaar, 1997; PerezGussinye et al., 2004; Poudjom Djomani et al., 1999; Rohrmann et al., 2002) all indicate that the Scandes have a lower flexural rigidity than the centre of the Fennoscandian Shield. Poudjom Djomani et al. (1999) and Rohrmann et al. (2002) conclude that the Scandes have low flexural rigidities of the order of 1.5–7.5 × 10 22 to 1 × 1023 Nm in the north and b1 × 1021 in the south. Ebbing and Olesen (2005) calculated a maximum flexural rigidity of 1 × 1023 Nm in the southern Scandes with decreasing values to the north. All these studies are based on the relation between the gravity signal (either free-air or Bouguer anomaly) and topography. These methods suffer from the absence of prominent topography over the central Fennoscandian Shield. McKenzie (2010) demonstrated the difficulties in obtaining realistic values for areas with subdued topography and under influence of recent uplift, factors which make an interpretation of the absolute values difficult. The pattern of regional compensation is however consistent and all the studies show that the lithosphere at least below the Scandes is rather elastic. The isostatic residual map (Fig. 5B) and the discussions of the geophysical data show that large deviations from simple Airy–Heiskanen isostatic equilibrium exist and other structures have to be considered. Consequently, the loading applied by the local

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crustal density distribution is important to consider in a discussion of the isostatic state of the Scandes. 4.3.2. A 3D isostatic model of the Scandes In the following, we present a model that explains the isostatic loading of the lithosphere assuming a combination of Pratt and Airy isostatic equilibrium, after Ebbing (2007) and which has been updated according to the results of the recent experiments and additional data. The model does not consider all the elements of the complex geological structure of the Scandes, but rather provides a model, which explains the gravity field to a large degree, assuming isostatic equilibrium. If we assume conventional Airy–Heiskanen isostasy the mass of the high topography would be compensated by a crustal root (Fig. 5A). The isostatic compensation occurs actually at the base of the lithosphere, but the crust–mantle boundary is associated with a high density contrast (~ 300 kg/m 3), which is far higher than the contrast usually assumed at the base of the lithosphere (~20–40 kg/m 3). Therefore, the crust–mantle boundary is the major contributor to the isostatic balance. However, over Fennoscandia, we observe from seismic studies a Moho depth which is deepest beneath the areas of low topography (compare Figs. 1 and 8). In this case, the lithosphere is not in isostatic equilibrium and the gravity field is not explained. Therefore, a high-density lower crust can be calculated that balances the isostatic state for the Scandes and its surroundings, by a combination of Airy and Pratt isostasy. We use the crustal base according to Fig. 8 and the base lithosphere (Fig. 11), and balance the model isostatically by a highdensity lower crust. Seismic studies (e.g. England and Ebbing, in press; Korsman et al., 1999) provide support for a minor density contrast (200 kg/m 3) between crust and mantle as the density contrast used in the Airy-isostatic calculations (see Section 3.2.3). The insufficient distribution of seismic lines does, however, prevent a clear definition of the extent of the high velocity lower crust from the seismic data alone. The densities of the isostatic model are applied according to the densities in Table 2. The geometry of the high-density lower crust is calculated by estimating the required loading to achieve isostatic equilibrium for the lithosphere as outlined in Ebbing (2007). The model is balanced only for the missing masses as the mass surplus cannot be related to the high-density lower crust. This approach leads to a thickness of the high-density lower crust as shown in Fig. 16A. The layer reaches thicknesses up to 25 km in the central Fennoscandian Shield, which is in agreement with estimates from seismic studies (e.g. Korsman et al., 1999). This high-density lower crust compensates for the deep Moho in the central Fennoscandian Shield and tapers out below the Scandes (Figs. 13–15 and 16A). The depth to the top of the high-density lower crust resembles a crustal root, and one can argue that a crustal root exists here (Fig. 16B). In some areas (e.g. beneath the Bothnian Sea) the high-density layer exceeds 25 km in thickness, which is certainly exaggerated. A reason for this might be that the density of the lower crust has to be even higher than applied in our model in areas where crustal thickness exceeds 50 km. Earlier density modelling indicates such high densities and no clear density contrast between the lower crust and upper mantle beneath the central Fennoscandian Shield

Table 2 Densities applied in the isostatic model. In the isostatic model, only the density contrasts control the isostatic response to the loads imposed by topography. Body

Relative densities [kg/m3]

Crust High-density lower crust Lithospheric mantle Asthenosphere

0 + 150 + 300 + 280

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Fig. 16. Isostatic high-density lower crust: A) Thickness of isostatic high-density lower crust. The calculated thickness is necessary to achieve isostatic equilibrium. B) Depth to isostatic high-density lower crust. A Gaussian 150 km lowpass-filter has been applied to the isostatic high-density lower crust grid to smooth high-frequency variations. Note: Only positive thickness values have been used to calculate depth to layer. Where high-density lower crust is thin or absent, the map represents depth to Moho. The orange lines indicate the location of profiles A, B, C in Figs. 13–15.

(e.g. Glaznev et al., 1996; Kozlovskaya et al., 2004). While Kukkonen et al. (2008) argue for an eclogitic origin of the high-velocity, highdensity lower crust the apparent correlation between its westward edge with the presence of the TIB, might support a magmatic origin similar to the Oslo Graben, and we cannot rule out a combination of magmatic underplating and eclogitisation. From the isostatic model, we can also calculate its gravity signal. For the gravity modelling, we add a three-layered crust (see Table 1), and the same deep geometry as for the isostatic model. However, we only apply the high-density lower crust, where it has a positive thickness, e.g. provides additional loading to achieve isostatic equilibrium. Areas of negative thickness indicate that lowdensity material is required by the isostatic model, e.g. in the upper mantle. An additional feature necessary to explain the gravity field and to a minor degree to balance the model isostatically is the Trans-Scandinavian Igneous Belt (TIB). The TIB extends beneath the Scandes from the east. A low-density body representing the TIB is introduced into the upper crust (2640 kg/m 3) with a thickness of 12 km and a lateral extent as defined in Ebbing (2007, his Fig. 7) and inferred from the magnetic anomaly map (Fig. 3). The gravity residual (Fig. 17) shows for most of the area residuals b ± 20 mGal, which is a good fit for such a general model. But the Northern and especially Southern Scandes are associated with gravity lows. Consequently, a low-density domain may be found at shallow depth below the Moho, if the crustal thickness estimates are correct. In the cross-sections (Figs. 13–15), we compare the depth to the top of the high-density lower crust with the detailed density models.

There is, in general, good correlation between the seismic and isostatic high-velocity, high-density lower crust, which adds weight to the choice of parameters for the isostatic model. In the Central Scandes (Fig. 14) the high-density lower crust correlates almost perfectly with the seismically derived high-velocity lower crust to the east of the Scandes. Towards the area of high topography, the isostatic lower crust continues further to the west before it tapers out and has a substantial thickness beneath the highest topography. Both the seismic and isostatic geometry of the high-density lower crust can be used to explain the gravity field along the Central Scandes, but with the seismic geometry of the middle and lower crust the profile is easier to balance against the offshore section. Comparison to the Southern Scandes makes it more likely that the high-density lower crust is only very thin or absent beneath the Scandes. For the Northern and Southern Scandes, the isostatic LCB has been used as an input in the cross-sections and hence correlates very well to the west of the Scandes. However, there are some interesting observations, where the high-density lower crust could not be applied to the model. In the Southern Scandes, the highdensity lower crust crosses below the seismically imaged Moho beneath areas of high topography. This means that isostatic support requires low-densities in the upper mantle. The areas coincide with the positive anomalies visible in the gravity residual (Figs. 5B, 7B and C). In the model for the Northern Scandes the area of deviation from isostasy is very small and within the uncertainties of the constraints, as the profile is located just south of the main residual anomaly.

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is less than the minimum wavelength over which typically isostatic compensation occurs. 5. Conclusions and outlook

Fig. 17. Gravity residual of the 3D isostatic model calculated using GMSYS-3D by subtracting the gravity effect of the model with the geometry of Figs. 8, 11 and 16B and the density contrasts in Table 1 from the Bouguer anomaly in Fig. 4B.

The general geometry of the crust and especially of the Moho is important for exact calculation of the isostatic response. For the Northern Scandes, the results of the SCANLIPS 2 experiments are needed to provide a better constraint on the crustal structure and the source of the gravity signal. In the Southern Scandes, the results by Svenningsen et al. (2007) reveal a crustal thickness 3–4 km thicker than the estimates from Stratford et al. (2009). Such a thick Airy-type crustal root beneath the highlands of southern Norway would leave little room for additional buoyancy-effects below the Moho. If we use the Moho depth of Svenningsen et al. (2007) in our isostatic model, the high-density lower crust would be very thin or absent beneath the Southern Scandes. In any case, it is clear that the highdensity lower crust tapers out at least below the Southern Scandes. Additionally, recent seismological studies point to a transition in upper mantle structure from Southern Norway across the Oslo Graben and into the central Fennoscandian Shield (e.g., Medhus et al., 2009, in press) with reduced shallow upper mantle velocities under the high topography. Another area where the modelled cross-section and the isostatic model deviate is the Oslo Graben. The presence of a high-velocity lower crust beneath the Oslo Graben has been confirmed by recent seismic studies (Stratford et al., 2009). The isostatic model shows a thinner high-density lower crust than the cross-sections. This might be related to the internal crustal structure of the crust in the Oslo Graben, which is more complex than the simple isostatic model. As mentioned in Section 4.2.1, comparison of receiver function data and wide-angle data may indicate a difference in the lithology of the high-density lower crust beneath the Oslo Graben and beneath the Fennoscandian Shield. Furthermore, the isostatic model does not consider the magmatic rocks of the Oslo Graben, which extend to 15 km in depth, but which have a horizontal width of b150 km, which

A key to the improved understanding of the mountain belt will be to link the detailed studies from the Southern (e.g. Stratford et al., 2009; Svenningsen et al., 2007) and Central Scandes (e.g. England and Ebbing, in press; Schmidt, 2000) to new installations for the Northern Scandes, to understand the differences in the lithospheric architecture and their impact on the observed topography. All available data show that increased topography correlates with an increase in crustal thickness relative to the Norwegian coast, but that the crust remains thick or even deepens towards the east. The pre-Caledonian structure of the crust east of the Scandes, especially the high-density lower crust, appears to dominate the isostatic state and explains the absence of a simple Airy-type crustal root. The high-density lower crust and the TIB reflect structures that appear to have only been slightly affected by the Caledonian orogeny and later processes. The location of the TIB coincides with the westward extension of the high-density lower crust, which might point to a causative link between the high-density lower crust and the emplacement of the TIB. This might indicate a mafic origin for the high-density lower crust. The data and models presented here show that an abundance of geophysical data is available over the Scandes but that gaps remain, especially for the Northern Scandes. The new results in Southern Norway and the Central Scandes have increased our knowledge of the lithospheric architecture and have confirmed the difference between the different segments of the Scandes mountain chain. Further integration of different types of geophysical data will hopefully help us to build a consistent model for the entire Scandes in the future. The presented data and models provide a basis to evaluate some of the proposed models for the Neogene uplift of the Scandes. For understanding the development of the topography of the Southern Scandes, the origin and age of the anomalous mantle beneath the Southern Scandes is of particular interest. Either the mantle contributed to a different buoyancy between the Southern and the Central and Northern Scandes since its emplacement during the Caledonian orogeny, or the mantle density has been decreased during more recent processes. The precise role of the differences in the lithospheric mantle beneath Southern Norway and Sweden is the topic of ongoing research efforts (e.g., Kolstrup, 2010). The transition in mantle properties cannot, yet, be followed as far north as the profile crossing the Central Scandes. Therefore, the continuation or disappearance of the transition zone is certainly of interest for understanding the lithospheric system of the Scandes. North of our Central Scandes profile, there is clearly a need for data acquisition with methods that are able to resolve the deep structure of the mountain belt and its surroundings. A refined model for the upper mantle can be expected from ongoing seismological studies from the MAGNUS data (Weidle et al., 2010). In addition results from the magnetotelluric experiments and the SCANLIPS 2 profile in Northern Norway will be available in the coming years and facilitate modelling of the processes that shape the topography of the Scandes. With respect to the shallow lithosphere, it will also be of importance to make further efforts to integrate near-surface data from Norway and Sweden. Integration of petrophysical data and near-surface seismic and MT data will enable us to link geology to depth for the entire Scandes. The extension of surface geology to depth is certainly of interest, as this will integrate geological observations with deep geophysical studies. This research is a focus of the ongoing TOPO-EUROPE initiative (e.g., Cloetingh, 2007), which aims to increase our understanding of the different process acting at different wavelengths in shaping surface topography.

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