New Moho Map for onshore southern Norway

Transcription

1 Geophys. J. Int. (2009) 178, doi: /j X x New Moho Map for onshore southern Norway Wanda Stratford, 1 Hans Thybo, 1 Jan Inge Faleide, 2 Odleiv Olesen 3 and Ari Tryggvason 4 1 Department of Geography and Geology, University of Copenhagen, Copenhagen, Denmark. Ws@geo.ku.dk 2 Department of Geosciences, University of Oslo, Oslo, Norway 3 Geological Survey of Norway, Trondheim, Norway 4 Department of Earth Sciences, Uppsala University, Uppsala, Sweden Accepted 2009 May 6. Received 2009 May 3; in original form 2008 August 6 1 INTRODUCTION Over the last 30 yr seismic studies of the lithosphere in southern Norway have led to models that predict a range of Moho depths beneath the Southern Scandes Mountains (Fig. 1b). Early studies relied on coarse refraction profiling to sample Moho depths (Sellevoll & Warrick 1971; Kanestrøm 1971; Tryti & Sellevoll 1977; Mykkeltveit 1980; Cassell et al. 1983) and interpolation between surveys to infer regional trends (Sellevoll & Warrick 1971; Kinck et al. 1993; Fig. 1b). Large receiver spacing and gaps in the coverage of these early studies meant that high resolution imaging of the Moho beneath the mountains was lacking. Results from early refraction profiling indicate that the crust is only 38 km thick (Sellevoll & Warrick 1971; Fig. 1b), and therefore, mechanisms other than variations in crustal thickness have been inferred for isostatic support of the observed topography (Olesen et al. 2002; Ebbing & Olesen 2005; Ebbing 2007). A Bouguer gravity anomaly of 80 to 130 mgal, centred on the Southern Scandes Mountains, is suggestive of a substantial density anomaly at depth. A correlation between topography and Bouguer gravity has been used to infer Airy-isostatic support for the mountains (Olesen et al. 2002; Ebbing & Olesen 2005) and the presence of a crustal root (Balling 1980). Recent Receiver Function studies (Fig. 1b) supported this argument as contrary to the SUMMARY A recent seismic refraction study across southern Norway has revealed that the up to 2469 m high Southern Scandes Mountains are not isostatically compensated by a thick crust. Rather, the Moho depths are close to average for continental crust with elevations of 1 km. Evidence from new seismic data indicate that beneath the highest topography Moho depths are around km. These measurements are 2 km deeper than early estimates interpolated from coarsely spaced refraction profiles, but up to 3 km shallower than Receiver Function estimates for the area. 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. However, the highest mountains do not overlie the thickest crust and additional support for topography, for example from flexural strength in the lithosphere, low densities in the upper-mantle or mantle dynamics, is likely. The relationship between topography and Moho depth breaks down for the Oslo Graben and the Fennoscandian Shield to the east and north. High density lower crustal rocks below Oslo Graben and increasing crust and lithospheric thicknesses below the Fennoscandian Shield may produce a negative correlation between topography and Moho depth. Key words: Controlled source seismology; Continental margins: divergent; Dynamics of lithosphere and mantle; Crustal structure. early refraction profiling, a crustal root and Moho depths of up to 43 km are interpreted beneath the mountains (Svenningsen et al. 2007). A negative Bouguer gravity anomaly, however, does not necessarily demonstrate the presence of a crustal root, as the depth distribution of densities is unknown and there can be dynamic support from the mantle. Bouguer anomalies of similar magnitude and sign have been observed in other mountain ranges with 1200 m mean elevations such as the Southern Alps in New Zealand (Koons 1994; Stern 1995) and the northern Transantarctic Mountains (ten Brink et al. 1997; Studinger et al. 2004; Stern et al. 2005). Despite the similar mean elevations and Bouguer gravity signatures, the Southern Alps have an over thickened crustal root and high mantle densities to balance the surface topography (Scherwath et al. 2003), whereas the Transantarctic Mountains are supported, at least in part, by a low density mantle (Stern & ten Brink 1989; Lawrence et al. 2006). For the Southern Scandes, flexural modelling (Ebbing & Olesen 2005) has been used to infer that the gravity anomaly is produced, at least in part, by low densities below the Moho. Presented here are new results from Magnus-Rex (Fig. 1a), a recent refraction seismic study across southern Norway. The study was undertaken to add required high resolution constraints to, and update, the Moho depth map for the Southern Scandes Mountains and to determine if the mountains are supported by a crustal root. GJI Tectonics and geodynamics C 2009 The Authors 1755

2 1756 W. Stratford et al. Figure 1. (a) Magnus-Rex (Mantle investigations of Norwegian uplift structure) is the refraction component of the MAGNUS project, a multidisciplinary study of the lithosphere in southern Norway. Magnus-Rex seismic experiment: Red stars are shot locations, blue dots are seismograph stations. Line numbers and shot numbers refer to diagrams in Figs 2 and 3. Insert: map of northern Europe showing the study area. (b) Contour map of Moho depth from Kinck et al. (1993). Contours show Moho depth in km. The map is based on a collection of early refraction studies in the area. Lines are as follows: (A) Fedje-Grimstad (Sellevoll & Warrick 1971). (B) Cannobe, South Norway (Cassell et al. 1983). (C) Oslo-Trondheim (Kanestrøm 1971). (D) Otta-Årsund (Mykkeltveit 1980). (E) Flora-Åsnes (Sellevoll & Warrick 1971). (F) Oslo Graben (Tryti & Sellevoll 1977). (G) Trondheim-Sundsvall (Vogel & Lund 1971). (H) A crustal scale refraction profile with Moho depths inferred from reflections (PmP; Iwasaki et al. 1994). (I) and (J) recent Receiver Function profiles across southern Norway (Svenningsen et al. 2007). Green shaded area is the region where the middle and lower Caledonide nappes are preserved in southern Norway. Pink shows the regions where upper allochthon rocks are still present. Blue shading outlines the Oslo Graben. 2 TECTONIC SETTING Bordered to the west by the passive margin of the North Atlantic and to the east by thick crust of the Fennoscandian Shield, the Southern Scandes Mountains have undergone a number of uplift and subsidence events (Lidmar-Bergstrøm et al. 2000; and references therein). Folding and faulting were induced by the collision between Baltica and Laurentia during the Caledonian orogeny from 440 to 410 Ma (Andersen 1998). During this period of mountain building, allochthonous nappes were over-thrust onto Fennoscandian basement rock from the west (Roberts & Gee 1985). The nappes can be divided into four main units, of which, three are present in southern Norway. Upper Allochthon Caledonides are preserved as far south as 62 N (Andersen 1998). A corridor of predominantly middle and lower allochthon Caledonides remain within the Faltungsgraben, a down-faulted graben associated with the postorogenic extensional collapse of the Caledonides in the southwest (Fig. 1b). The middle and lower allochthons represent the nappe sequences formed by collision of the continental margin of Baltica. In contrast, the upper allochthon is comprised predominantly of sedimentary and igneous rocks derived from the pre-collision crust of the Iapetus ocean, and include ophiolites and island arc complexes (Stephens 1988). Outside the Caledonides, basement rock in southern Norway is exposed and is predominantly comprised of granites deformed in the Sveconorwegian orogeny (Koistinen et al. 2001). Following the Caledonian orogeny, there was a reversal in deformation polarity and the Scandes underwent extension and postorogenic collapse (Andersen 1998). Despite this extension, a long lived mountain range throughout the Palaeozoic and Mesozoic has been argued (Ziegler 1988). For the Cenozoic, a number of phases of uplift have been inferred (Anell et al. 2009). The main phases being: uplift due to incisional erosion and other unknown mechanisms in the Neogene (Riis 1996; Lidmar-Bergstrøm et al. 2000; Faleide et al. 2002) and recent uplift of 1 4 mm yr 1 due to postglacial rebound from removal of the Fennoscandian ice sheet (Niskanen 1939; Balling 1980). Although postglacial rebound and incisional erosion are still in effect for the present day surface dynamics (Riis & Fjeldskaar 1992), other tectonic processes may be required to explain the inferred longer term uplift. 3 PREVIOUS STUDIES Refraction profiling began in earnest in the southern Norway region in the early 1970s. These studies used small numbers of analogue instruments that were relocated between shots to obtain the long offsets required to record Pn arrivals (Sellevoll & Warrick 1971; Mykkeltveit 1980; Cassell et al. 1983). The velocity models derived for the crust from these early studies are generally of two or three homogenous layers. Nevertheless, these early profiles show consistency between the observed velocity structures with depth throughout the southern Norway region. From these studies, velocities for the upper crust in southern Norway are of 6.25 ± 0.1 km s 1 and for the lower crust of 6.75 ± 0.05 km s 1 (Fig. 2). Although the near surface velocities differ between surveys, there is good agreement between crustal velocity profiles below 5 km Journal compilation C 2009RAS

3 Moho map, southern Norway 1757 Figure 2. Profiles of 1-D velocity structure with depth extracted from previous seismic velocity models for the crust in southern Norway (Fig. 1b). The locations and references for these velocity profiles are shown in Fig. 1(b) and the points where the 1-D velocity structures were extracted are marked with black dots. A 1-D profile extracted from line 1, this study, is also shown (location is marked with a white dot in Fig. 1a). depth. The interpretation of the Moho depth from these studies was in part questioned due to the coarse resolution of these early models. With large receiver spacings, these surveys were susceptible to the hidden layer problem whereby arrivals from thin layers at depth are not observed as first arrivals and depths can be over or underestimated (Kearey et al. 2002). Gaps between the surveys also left regions where significant extrapolation was required to produce the original maps of contoured crustal thickness (Sellevoll & Warrick 1971; Kinck et al. 1993; Korsman et al. 1999). 4 DATA AND RESULTS Magnus-Rex, a new crustal scale seismic refraction experiment, was conducted in October Three 400 km long seismic lines were deployed across southern Norway (Fig. 1a). The elevated central Southern Scandes were targeted with two profiles in a large X- pattern. An additional profile to the south extends east west across the Southern Scandes, through the Oslo Graben and into Sweden (Fig. 1a). A total of 26 shots of kg charge size were fired along the three lines. End shots were 400 kg. Smaller charges of 100 and 200 kg were distributed along the profiles. Ca. 750 vertical component (Texan) seismographs were deployed at two km spacing along the lines. A close instrument spacing of 750 m was used in a 120 km section across the Oslo Graben. Processing of the seismic data included basic frequency filtering (bandpass 2 4 to Hz) and trace balancing. The upper crust is well constrained by first arrival (Pg) refractions from all shots. No distinct changes in slope on the Pg arrivals branch are observed (Fig. 3); rather there is a gradual increase in slope with offset indicating an almost continuous increase in velocity with depth. Apparent velocities are 5.8 km s 1 at near offsets grading to 6.6 ± 0.1 km s 1 at offsets >150 km (Fig. 3). Velocities in the middle crust are constrained by first arrival refractions beneath the centre of the lines. The lower crust and the layer of high velocity (>7kms 1 ) material at the base of the crust below the Oslo Graben are constrained by the moveout of reversed PmP (Moho) reflections. Pn arrivals from beneath the centre of all three lines have apparent velocities of 8.05 ± 0.1 km s 1 in all directions (Fig. 3) and crossover distances with Pg of km. The accuracy of picking seismic data depends on the noise level and frequency of the arrivals and for crustal refraction data, arrivals can normally be picked to within a half wavelength. For the Magnus- Rex data, uncertainties in pick arrival times of 0.1 s are estimated. Based on regression analysis of these traveltime picks, velocities are estimated to have errors of ±0.05 km s 1 for Pg phases and ±0.1 km s 1 for Pn. Forward modelling ray tracing using RAYINVR (Zelt & Smith 1992) is undertaken to model the observed traveltimes of Pg, PmP and Pn. A 2-D velocity model is constructed from the observed traveltimes by projecting all arrivals onto a 2-D plane by conservation of shot offset (Zelt & Smith 1992). Seismic models are constructed by a top down approach whereby near surface velocities and velocity structures are modelled first. Data-dependent errors are estimated from the root mean square (rms) misfit of the model to the traveltime picks, which is optimally within the pick uncertainties (Table 1). The normalized χ 2 or chisquared statistic represents how well the model traveltimes fit the observed traveltimes to within the pick uncertainties χ 2 = 1 n 1 n [ ] (t(i)calc t(i) obs ) 2, σ (i) i=1 where t(i) calc and t(i) obs are the ith calculated and observed traveltimes, σ (i) is the traveltime uncertainty and n is the number of observations. Values less than 1 indicate an overfit of the model to the data and X 2 values of 1 represent the best obtainable fit. Model dependant uncertainties are estimated by forcing the ray tracing solutions into end-member velocity models. The upper crust and middle crust contribute the least to the overall uncertainty in Moho depth as there is complete, reversed ray coverage from all shots. However, there is no first arrival coverage of the high-velocity lower crust. Thus, bounds on the depth to the Moho are estimated by the degree to which the structure of this high-velocity layer can be altered, while still retaining a fit to the observed traveltimes within the pick uncertainty (Table 1). Where there is coverage of Pg, PmP and Pn the uncertainty in Moho depth is ±1 km. Where constraint is from Pg and PmP only, the uncertainty is ±2 km. In refraction profiling hidden layers can cause errors in depth determinations. However, the joint analysis of wide-angle reflection and refraction data can limit the range of possible hidden layer thicknesses and velocities (Sain & Reddy 1997). For southern Norway, appropriate end-member velocities for possible low- and highvelocity layers are inferred to be 5.4 (e.g. metagreywacke) and 7.4 km s 1 (gabbro; Christensen & Mooney 1995; Koistinen et al. 2001). Applying these velocities, hidden high- and low-velocity layers in the crust up to 3 km thickness produce Moho depths within the

4 1758 W. Stratford et al. Figure 3. Examples of seismic data from the Magnus-Rex Project. Data are plotted at reduced traveltime where T = t offset/8. (a) Shot 1, Line 1: Crustal refraction (Pg) P-waves velocities (V p ) of km s 1, Moho reflections (PmP), and arrivals from below the Moho, a mantle reflection and refracted arrivals (Pn) at offsets of 160 km, and with V p of 8.05 km s 1. (b) Shot 18, Line 2: Pg ( km s 1 ), PmP and crustal reflectivity (PiP), Pn arrivals at offsets of 180 km with apparent V p of 8.05 km s 1. (c) Shot 26, Line 3: Pg ( km s 1 ), PmP reflections, Pn arrivals at offsets of 180 km with apparent V p of 8.05 km s 1 and a strong reflection from the mantle at offsets of km. uncertainties bounds. Moreover, no traveltime skips can be seen in the first arrival data that might elucidate a hidden layer of significant thickness. 5 CRUSTAL VELOCITY STRUCTURE Although crustal velocities beneath the Magnus-Rex lines appear to be almost laterally homogeneous, there are variations in the top 5 km of the crust (Fig. 4). This may be, in part, due to the high ray coverage in this layer. However, ray coverage in the top 10 km is adequate to highlight significant lateral velocity variations. Therefore, we attribute the observed velocity variations in the top 5 km to the Caledonide nappe sequences on lines one and two, and on line three to the volcanic rocks and sediments within the Oslo Graben. Near surface velocities along line 1 decrease from 6kms 1 in the northwest to 5.8 km s 1 in the southeast (Fig. 4a). Although it is not possible to attribute these velocity changes to specific Journal compilation C 2009RAS

5 Table 1. Assigned pick uncertainties and calculated model misfits (rms and χ 2 values): n is number of observations, σ is assigned pick uncertainty, t rms is the rms misfit of the model to the picks, and X 2 is the normalized chi-squared value. n σ (s) t rms (s) X 2 Line 1 Pg PmP Pn Line 2 Pg PmP Pn Line 3 Pg PmP Pn composition changes, the profile line crosses out of Sveconorwegian basement rock (predominantly granite and granodiorite) in the northwest into the middle and lower allochthon of the Caledonian nappe sequences in the southeast. Near surface velocities along line 2 decrease from 6 kms 1 in the south to 5.8 km s 1 in the north (Fig. 4b). The profile line tracks along the western edge of the nappe sequence in the southwest and near the crossover point with line 1, passes into the upper allochthon in the north. The high velocity body with V p of 6.5 km s 1 (at distances of km on line 2) at a depth of 3 km is possibly ophiolite rock associated with the upper allochthon (Roberts & Gee 1985; Andersen 1998). This high velocity body is observed on shot 11 arrivals only, thus its lateral and depth extent are not well constrained. Previous seismic surveying in this area (line D on Fig. 1b) also found anomalous velocities at depth beneath the nappe sequences in this region. A 4 km thick low-velocity body at 14 km depth has been inferred from an offset in first arrival refraction on an array just to the southwest of shot 11 (Mykkeltveit 1980). Anomalous velocity-depth relationships may, thus, be a feature of this region of the Caledonides, although the different depth ranges of the putative bodies indicates that the variation may be attributed to more than one allochthon. Near surface velocities on line 3 show the most variation in the Oslo Graben (Fig. 4c). Velocities as low as 5.5 km s 1 here are attributed to Cambro-Silurian sediments and Permo-Carboniferous granitic rocks within the graben (Neumann et al. 1992). In the west near surface velocities are 5.8 km s 1 where the line crosses the middle and lower allochthons. Between distances of 100 and 240 km the near surface velocities are around 6.0 km s 1 within the Sveconorwegian basement. Velocities in the upper to middle crust are more homogeneous along all three lines with a general velocity increase with depth, grading from 6.1 to 6.4 km s 1 at around 22 km depth. The notable exception is the Oslo Graben where higher velocities of 6.6 km s 1 are inferred to occur at depths as shallow as 10 km. These high velocities are focused toward the western side of the graben (Fig. 4c). The lower crust is constrained by refracted arrivals on all three lines with velocities of 6.6 km s 1 at a depth of 22 km except for the gap due to the Oslo Graben, where velocities at this depth are less well constrained (Fig. 4c). The lower-most crust where velocities of km s 1 are inferred immediately above the Moho is not constrained by refracted arrivals. Instead the thickness and velocity Moho map, southern Norway 1759 of this layer has been determined from the best-fit model to explain the moveout of reversed PmP reflections from the Moho. Variations in V p and thickness of this layer of about ±0.05 km s 1 and ±1 km, respectively, are possible within the uncertainties of the PmP picks. Beneath the highest mountains a Moho depth of 40 ± 1km is modelled (Fig. 4a). A Moho depth of 39 ± 1 km is matched on lines 1 and 2 at their cross point (Figs 4a and b). Constraint on Moho depth extends to the southwest of the crossing point on Line 2 and indicates a decrease in Moho depth to a flat section at 38 ± 1 km at distances of km and 36 ± 2kmbyadistanceof 110 km from the south end of the line (Figs 4b). This result concurs with an earlier reflection profile along Sognefjord that led to the interpretation of 31 km thick crust near the coast and 36 km thick crust by 100 km inland (100 km from the west coast along profile H, Fig. 1b; Iwasaki et al. 1994). There is also consistency with results from two regional refraction profiles recorded in 1971 (Sellevoll & Warrick 1971) where depths of 38 km were inferred just north of Sognefjord (160 km distance along profile E, Fig. 1b) and 36 km at 150 km southeast of Bergen (150 km distance along profile A, Fig. 1b). Along Line 1, the Moho depth decreases to around 38 ± 1 km depth just to the northwest and southeast of the cross point between the lines 1 and 2 (Fig. 4a). As the Oslo Graben is relatively narrow, the profile length in the graben is insufficient for first arrivals from the lower crust to be resolved. Thus, less constraint for lower crustal structure is available from Line 3 and results from a previous north south oriented seismic line (Tryti & Sellevoll 1977; profile F, Fig. 1b), are therefore, included in the velocity modelling (Fig. 4c). Tryti & Sellevol (1977) inferred velocities of 7.1 km s 1 at just over 20 km depth in the graben. Using this added constraint, the Moho up-warp beneath the graben is interpreted to be small, of the order of 2 km, giving a Moho depth of 34 ± 2 km, which is in accord with the interpretation of Tryti & Sellevol (1977). West of the Oslo Graben the Moho depth is 36 ± 1 km at 180 km and 38 ± 1 km by 140 km from the west end of the profile line (Fig. 4c). Earlier studies of the Graben interpreted shallower Moho depths of from reflection (Kanestrøm & Haugland 1971) and gravity modelling (Ramberg & Smithson 1971). However, from the late 1970s there have been a number of studies that indicate a smaller Moho up-warp beneath the Graben (Ramberg 1976; Tryti & Sellevoll 1977), which is supported by the new data presented here. A smaller Moho up-warp is similar to findings of little or no Moho topography at currently active rifts (Baikal Rift and Kenya Rift; (Thybo et al. 2000; Nielsen & Thybo 2009; Thybo & Nielsen 2009) and at the DonBas Graben in Ukraine (Lyngsie et al. 2007). Thybo & Nielsen (2009) explain these observations by the process of magma-compensated crustal thinning, whereby intruding magma compensates for the crustal thinning taking place due to extension and stretching. 6 NEW MOHO MAP FOR SOUTHERN NORWAY Moho depth measurements from Pn and PmP arrivals are plotted at their geographic locations and contour lines are hand drawn between the values (Fig. 5). The simplest means of extrapolation are used and the contouring process aims to limit extra structure. However, there are still gap regions within the data set. Under regions where seismic constraints from Pn arrivals from this study are applied, the Moho depth map has uncertainties of ±1 km, where PmP arrivals only available the uncertainty is ±2 km. Where Moho depth constraints

6 1760 W. Stratford et al. Line 2 a.) Moho b.) c.) Velocities in km/s Velocities in km/s West East Velocities in km/s Velocities in km/s South Line (2) Oslo Graben Velocities in km/s Line 1 Moho Line 2 North Moho Line 3 West East Figure 4. Forward modelling ray tracing solution for lines 1, 2 and 3. See Fig. 1 for line locations. The forward modelling solution uses a top down approach to fitting picked arrival times. Solid lines indicate where constraint is available from refractions. Thick dashed lines indicate where constraint is from PmP reflections only. Thin dashed lines are the velocity layers boundaries. A discussion of model uncertainties is given in the text. The number of traveltime picks, and the pick and model uncertainties are give in Table 1. The cross point between lines 1 and 2 is marked on the seismic models with labelled arrows. The surface expression of the Oslo Graben and dotted lines demarking the high velocity 6.6 km s 1 rock in the graben are marked on line 3. from earlier refraction studies are used (seismic profiles shown in Fig. 1b), uncertainty in Moho depth is more difficult to quantify as uncertainties are often not assessed in early papers. However, the methods and data used indicate uncertainties of ±2 km would be adequate and these are in line with those inferred from PmP constrained Moho depths in this study. Overall, the new compilation map of Moho depths for onshore southern Norway is similar to that of Kinck et al. (1993), but the Journal compilation C 2009RAS

7 Figure 5. (a) New Moho map for southern Norway. Labelled contours represent Moho depth in km. The map contains most of the original depth measurements used in the map of (Kinck et al. 1993; Fig. 1b), as well as new measurements from this study. Singular black dots on the map are locations of depth measurements from early studies (see Fig. 1b for references and Fig. 2 for velocity information from these lines). Lines of dots are the seismic profiles of Magnus-Rex. Black dots are the region of Pn Moho depth constraint; Grey dots are where PmP depth constraints are available. The contour lines are superimposed on a map of topography that has been smoothed by averaging elevations over a 50 km window. Note there is some correlation between topography and Moho depth beneath to the Southern Scandes; however the highest mountains are offset to the west of the thickest crust. extra constraints from this study have added details and improved the reliability significantly. The offshore data included in Kinck et al. (1993) Moho map in Fig. 1(b), are not included in this plot as more recent studies using expanding spread profiles (ESP) of the Moho depth off the west coast provide better constraint on the crustal structure here (Christiansson et al. 2000). From Christiansson et al. (2000), seismic study a high velocity body in the lower crust near the coast was inferred and interpreted as eclogized rock (Christiansson et al. 2000). Eclogites are found at the surface in western Norway (Andersen 1998), however whether the putative high velocity body of eclogite extends onshore into the lower crust cannot be constrained by the new data presented here. Planned onshore offshore profiling may update the gradient in crustal thickness between onshore and the continental slope. New constraints provided by this study cover some of the gaps between the earlier refraction profile studies in southern Norway. The original Moho map remains unchanged along the lines of the prior studies and where there is overlap between the prior studies and the new data presented here, there is agreement on Moho depth within the uncertainties. The main difference is an extension of the slightly thicker crust (38 km) southward under the Southern Scandes Mountains (compare the 38 km contour in Fig. 1b with Fig. 5). Moho map, southern Norway 1761 A recent Receiver Function study along a profile line of similar coverage to the Magnus-Rex Line 1 has been used to infer the presence of a crustal root beneath the Southern Scandes (Svenningsen et al. 2007; profile I, Fig. 1b). A Moho depth of 38 km at the coast, thickening to 43 km beneath the centre of the line was interpreted. The general shape of the Moho interface inferred by the Receiver Function study concurs with that determined from the new active source refraction results presented here; but the depth to the interface from refraction profiling is around 3 km less. The differences in inferred Moho depth may, in part, be due to the assumed P-wave velocity model for the lower crust used in the Receiver Function depth migration being 3 per cent higher than what has been found from refraction profiling. This difference in velocity, however, can explain only 200 m of the depth difference between the two techniques. The remaining difference is at the extreme bounds of the combined uncertainties (±2 km) of the Receiver Function study and (±1 km) of the refraction profiling. Receiver Function techniques utilise the travel times of waves converted from P to S at the Moho and there is an inherent trade off between velocity and depth. Moreover, the trade off between velocity and depth is particularly strong for V S compared to V P and changes in V p /V S of only 0.1 can lead to 4 km change in crustal thickness (Zhu & Kanamori, 2000). The earthquake waves used in a Receiver Function study are also of significantly lower frequency, and hence lower resolution, than those frequencies used in active source refraction studies. 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. Therefore, we favour the shallower Moho depths determined by refraction profiling in this study. 7 AIRY-ISOSTASY We assess the degree of crustal support for topography by comparing the relationship between topography and crustal thickness to a first-order model of crustal buoyancy (isostatic Moho, Fig. 6). Here crustal thickness is taken as the depth to the refraction Moho. Assuming Airy-isostasy and complete compensation for topography by the crust, total crustal thickness (h t ) is related to topography (h c ) by ) ρ t h t = h i + h c (1 +, ρ m ρ c where h i is the crustal thickness at sea level for southern Norway (30 km; Kinck et al. 1993), ρ c (crust) and ρ t (topography) are the average density of the crust (2830 kg m 3, estimated from the average crustal velocity from this study and an empirical relationship between seismic velocity and density from Brocher (2005) using the original data of Ludwig et al. (1970)) and ρ m is the density of mantle lithosphere (3300 kg m 3 ). We use smoothed topography, with a 50 km diameter for the smoothing window, for h c. This formula assumes that the vertical density distribution in the crust is the same everywhere along the profile. An average crustal density is often used for isostatic calculations and for southern Norway gives a density contrast at the Moho of 470 kg m 3, which is larger than can be inferred from the lower crustal and upper-mantle seismic velocities in this study. Moreover, if the compensation is assumed to be at the crust mantle boundary then the density contrast at this boundary will be the compensating factor. Thus, two models are used. One of average crustal density (ρ c and ρ t = 2830 kg m 3 ) and one where topography (ρ t = 2670 kg m 3 ) is balanced by higher densities (ρ c = 2950 kg m 3 )

8 1762 W. Stratford et al. a.) Crustal thickness (km) West Line 1 Refraction Moho Isostatic Moho Isostatic Moho (LC) East Elevation (km) Distance (km) b.) South North 1.5 Line Refraction Moho Isostatic Moho 32 Isostatic Moho (LC) Crustal thickness (km) Crustal thickness (km) c.) West Line 3 Distance (km) Refraction Moho Isostatic Moho Isostatic Moho (LC) Distance (km) Figure 6. Graphs showing the Airy-Isostatic relationship between topography and Moho depth for southern Norway and south western Sweden. Moho depths and topography are extracted from Fig. 5 along the seismic profiles. Moho depths from the map (Fig. 5) and predicted from the topography/moho depth relationship based on two Airy-isostatic balance models are shown. Isostatic Moho is the model using an average crustal density to model topography and Isostatic Moho (LC) refers to the model where a higher density lower crust (ρ c = 2950 kg m 3 ) is used. Note the distinct differences in the relationship between the Moho depth and smoothed topography for the Southern Scandes, and for the Oslo Graben (distances km) and Swedish Fennoscandian crust (distances km). in the lower crust (below the 30 km depth of the reference model). The second model gives a density contrast of 350 kg m 3 across the Moho for southern Norway, that is in accord with the contrast inferred from seismic velocities near the Moho from this study. Both models are presented in Fig. 6 to illustrate the uncertainties in assuming densities with isostatic modelling. Profiles comparing refraction Moho depth (from Fig. 5) and the Moho depth predicted from the Airy-isostatic balancing of smoothed topography (eq. 1) are used to define the relationship between topography and crustal thickness in southern Norway. 2-D profiles (Fig. 6) are extracted along the three Magnus-Rex seismic lines from the contours of the refraction Moho map and the East smoothed topography. This topography is smoothed with a 50 km diameter filter to remove short wavelengths; smoothing out the peaks and the deeply incised, but narrow, fiords. Wavelengths of 50 km are significantly smaller than the topographic wavelengths that are likely to be locally compensated. There are uncertainties in the conversion of seismic velocities to densities and it is the pattern of the calculated isostatic compensation including form, wavelength and magnitude, which contains useful information on how topography is supported. Along line 1, the Airy-isostatic models have a maximum crustal thickness of 39 km, which is close to the 40 km interpreted for the refraction crustal thickness (Fig. 6a). However, the region of highest Elevation (km) 1.5 Elevation (km) Journal compilation C 2009RAS

9 elevation is not located above the thickest crust, but rather is offset 60 km to the west over a gradient in crustal thickness. An offset between the highest elevation and the thickest crust is observed at a number of mountain ranges and has been attributed to the load of the mountains being partially supported by the flexural strength of the lithosphere (Forsyth 1985; Stewart & Watts 1997). East of the highest mountains the crustal thicknesses are higher than those inferred from Airy-isostatic balance of topography (Fig. 6a). Line 2 is close to being in isostatic equilibrium for the southwest part of the line, but again the highest elevations (around distances of 180 km) remain under-compensated. The fit is poor at the northeastern end of the line where the crust thickens to values higher than the topographic elevations would suggest (Fig. 6b). Along line 3 the deepest isostatic Moho is also offset to the west of the deepest refraction Moho beneath the mountains (Fig. 6c). To the east the low elevations are over-compensated by a thicker crust than the isostatic Moho predicts. This is especially true for the Oslo Graben where evidence from higher seismic velocities at shallower depths in the graben indicate that the crust here is of higher density, which has not been accounted for with the current calculations. The low topography of Sweden is also over-compensated with the density structure of the isostasy models presented here. The indicated over-compensation in the Fennoscandian Shield proper may also be explained by the presence of a high density lower crust (BABEL Working Group 1993; Abramovitz et al. 1998) and a thicker lithosphere (Thybo 2001; Kaban 2003; Artemieva 2007; Artemieva & Thybo 2008). 8 DISCUSSION Despite the high elevations of the Southern Scandes, the peaks are bounded by deeply incised valleys such that the mean elevations, when averaged over 50 km 2, are only m (Fig. 5). The incised crust is slightly thinner than shield crust and relatively thin for an orogenic belt (Christensen & Mooney 1995). Simple 1-D isostatic calculations for southern Norway highlight the difficulty in explaining topographic elevations in regions where there are lateral variations in crustal structure and elastic thickness. For southern Norway, the principal evidence for deviation from Airy-isostasy (Fig. 6) is that the thickest crust is not located below the highest part of the southern Scandes Mountains. Based on the assumed density model, the Oslo Graben and the Fennoscandian Shield show no correlation with the models for Airyisostatic balance (Fig. 6). It is likely that the crust beneath the Oslo Graben may be of higher density than elsewhere in southern Norway (Ramberg 1976; Tryti & Sellevoll 1977; Ebbing et al. 2005). Furthermore, crustal thicknesses are high (>40 km) in the Swedish part of the Fennoscandian Shield where topography has low relief (<500 m; Korja et al. 1993; Fig. 5). A negative correlation between topography and Moho depth in the Oslo Graben and Fennoscandian Shield is observed (Fig. 6). The Fennoscandian Shield or parts of it may be underlain by thick lithosphere (Calcagnile 1982) with low density (Thybo 2001; Kaban 2003; Artemieva 2007; Artemieva & Thybo 2008) that will have an effect on lithospheric buoyancy and, therefore, on topography. For Airy-isostatic balance, however, the variation in thickness of the high velocity and hence high density, lower crust in Fennoscandia needs to be considered (Korja et al. 1993). A compilation map produced from seismic studies in Fennoscandia has inferred that the high velocity (>7 kms 1 )layer at the base of the crust varies between 0 and 25 km thick. Korja Moho map, southern Norway 1763 et al. s (1993) compilation map shows a general increase in the thickness of this layer to the north and east of the southern Scandes, with a thickness of 4 12 km beneath the mountains. However, only lines C and B (Fig. 1b) east of the mountains in southern Norway were used in this compilation and the contours were extrapolated to the west. A thinner high velocity lower crustal layer (<5 kmthick) is inferred from this study and this is a significant contrast to the predominantly 8 12 km thickness inferred for northern Norway and Sweden (Korja et al. 1993). Where the high velocity (high density) lower crust is thicker a higher average crustal density should be used to isostatically balance topography. Coherence techniques and gravity modelling have inferred effective elastic thicknesses (T e ) of 8 20 km for the Caledonides in southern Norway (Poudjom Djomani et al. 1999; Perez-Gussinye et al. 2004; Ebbing & Olesen 2005) increasing to 40 km in central Sweden (Perez-Gussinye et al. 2004). T e values beneath the southern Scandes are low to average for continental areas (Watts 2001). However, some degree of additional support from a lithosphere with elastic strength is inferred (Ebbing & Olesen 2005). A combination of crustal thickness and some flexural support may, thus, explain a significant component of the topography in southern Norway. However, better understanding of the elastic thickness variations across southern Norway is required before the magnitude of any additional topographic support mechanisms, such as additional positive buoyancy from low mantle densities, can be quantified. Thus for the highest peaks of the southern Scandes Mountains, full support of topography by crustal buoyancy cannot be demonstrated and density differences within the upper mantle may also be contributing (Olesen et al. 2002; Ebbing & Olesen 2005; Ebbing 2007). Further evidence for buoyancy within the upper mantle comes from earthquake Pn and Sn velocities. A zone of sub-moho low velocities has been inferred to extend beneath the Southern Scandes (Bannister et al. 1991). Similar low-velocity mantle has been inferred beneath other mountain ranges where crustal thicknesses are inadequate to explain elevations such as the Transantarctic Mountains (Lawrence et al. 2006) and the southern Sierra Nevada (Jones et al. 1994). Furthermore, effects on topography from dynamic processes must be considered. Recent studies using GPS and precise levelling measurements have inferred uplift values of 0 3 mm per yr for the southern Scandes (Danielsen 2001; Vestøl 2006; Lidberg et al. 2006). Although these rates decrease towards the southwest in the southern Scandes, a continued contribution from the effects of ice sheet removal may be in effect (Milne et al. 2004). Other possible mantle sources for dynamic topography have been inferred from seismological studies (Weidle & Maupin 2008). These studies of mantle S and P-wave speeds have inferred a low speed zone in the upper mantle below depths of around 70 km (Weidle & Maupin 2008). This observation, combined with the earlier sub-moho Pn and Sn study (Bannister et al. 1991), indicates that the mantle is anomalous beneath southern Norway. Such low mantle wave-speeds can be attributed to thermal or compositional anomalies and, by applying thermal expansion (Turcotte & Schubert 1982), a low density mantle can be inferred. A recent study combining gravity and heat flow data of the Norwegian mountains, however, showed that there are no visible effects of a mantle temperature anomaly at the surface (Pascal & Olesen 2009). What is causing the inferred mantle wave speed anomaly and what effect it has on mantle buoyancy remains a critical question for the magnitude and timing of the uplift and of the southern Scandes.

10 1764 W. Stratford et al. New data presented here outline the crustal contribution to topography in southern Norway. A region of deep Moho (38 40 km) extends north south beneath the Southern Scandes forming an apparent root between the extended crust of the Oslo Graben to the east and the Atlantic Margin/North Sea to the west (Figs 1a and 5). The data indicate that the Moho is slightly deeper beneath the central high mountains, than elsewhere in southern Norway. A deeper Moho indicates that more isostatic support is possible from the crust than previously inferred (Ebbing 2007). However, it is unclear how far the relationship between topography and Moho depth might extend. Flexural rigidity (Riis & Fjeldskaar 1992; Ebbing & Olesen 2005) and the increasing lithospheric and crustal thickness to the east and north beneath Fennoscandia (Calcagnile 1982) must be taken into account for assessments of isostatic balance of topography. Furthermore, the Southern Scandes are located at the transition between the Fennoscandian Shield and the North Atlantic Ocean, which may cause dynamic forces from edge driven convection in the mantle (King & Anderson 1998) or from release of compressive stresses associated with ridge push from the ocean into the shield (Thybo et al. 2000a). Regardless, the new Moho map presented here provides for the first time firm evidence on the relation between crustal thickness and the present elevation of the Southern Scandes Mountains. ACKNOWLEDGMENTS The instruments were provided by the PASSCAL facility of the Incorporated Research Institutions for Seismology (IRIS), PASSCAL Instrument Centre, New Mexico Tech, Socorro and the University of Copenhagen. Data collected during this experiment will be available through the IRIS Data Management Centre. Data acquisition was supported by the Danish National Science Research Council, the Carlsberg Foundation, the Norwegian Geological Survey and the Norwegian Research Council. Appreciation is extended to the principal field technicians, G. Kaip, B. Greschke, P. Jørgensen, J. Gellein, L. Furuhaug, G. Storrø, T. Sørdal and A. K. Nilsen. Thanks also to students from Copenhagen, Oslo, and Bergen universities and additional personal from the Norwegian Geological Survey who helped with the field deployment. Appreciation is also extended to J. Ebbing and R. England for constructive reviews of the paper. REFERENCES Abramovitz, T., Thybo, H. & MONALISA Working Group, Seismic structure across the Caledonian Deformation Front along MONA LISA profile 1 in the southeastern North Sea, Tectonophysics, 288, Andersen, T.B., Extensional tectonics in the Caledonides of southern Norway, and overview, Tectonophysics, 285, Anell, I., Thybo, H. & Artemieva, I.M., Cenozoic uplift and subsidence in the North Atlantic region: geological evidence revisited, Tectonophysics, in press, doi: /j.tecto Artemieva, I.M., Dynamic topography of the East European Craton: shedding light upon the lithospheric structure, composition and mantle dynamics, Global Planet. Change, 58, Artemieva, I.M. & Thybo, H., Deep Norden: highlights of the lithospheric structure of Northern Europe, Iceland, and Greenland, Episodes, 31, BABEL Working Group, Deep seismic reflection/refraction interpretation of critical structure along BABEL profiles A and B in the southern Baltic Sea, Geophys. J. Int., 112, Balling, N., The land uplift of Fennoscandia, gravity field anomalies and isostasy. in Earth Rheology, Isostasy and Eustasy, pp , ed. Mørner, N.A., John Wiley & Sons, New York. Bannister, S.C., Ruud, B.O. & Husebye, E.S., Tomographic estimates of sub-moho seismic velocities in Fennoscandia and structural implications, Tectonophysics, 189, Brocher, T.M., Empirical relations between elastic wavespeeds and density in the earth s crust, Bull. seism. Soc. Am., 95, Calcagnile, G., The lithosphere asthenosphere system in Fennoscandia, Tectonophysics, 90, Cassell, B.R., Mykkeltveit, S., Kanestrøm, R. & Husebye, E.S., A North Sea southern Norway seismic crustal profile, Geophys J. R. astr. Soc., 72, Christensen, N.I. & Mooney, W.D., Seismic velocity structure and composition of the continental crust: a global review, J. geophys. Res., 100, Christiansson, P., Faleide, J.I. & Berge, A.M., Crustal structure in the North Sea: an integrated geophysical study, in Dynamics of the Norwegian Margin, Vol. 167, pp , ed. Nøttvedt, A., Geol. Soc. Spec. Pub., London. Danielsen, J., A land uplift map of Fennoscandia, Surv. Rev., 36, Ebbing, J., Isostatic density modelling explains the missing root of the Scandes, Nor. J. Geol., 87, Ebbing, J. & Olesen, O., The Northern and Southern Scandes structural differences revealed by an analysis of gravity anomalies, the geoid and regional isostasy, Tectonophysics, 411, Ebbing, J., Afework, Y., Olesen, O. & Nordgulen, Ø., Is there evidence for magmatic underplating beneath the Oslo Rift?, Terra Nova, 17, Faleide, J.I., Kyrkjebø, R., Kjennerud, T., Gabrielsen, R.H., Jordt, H., Fanavoll, S. & Bjerke, M.D., Tectonic impact on sedimentary processes during the Cainozoic evolution of the northern North Sea and surrounding areas, Vol. 196, Geol. Soc. London Spec. Publication. Forsyth, D.W., Subsurface loading and estimates of the flexural rigidity of continental lithosphere, J. geophys. Res., 90, Iwasaki, T., Sellevoll, M.A., Kanazawa, T.V. & Shimamura, H., Seismic refraction crustal study along the Sognefjord, south-west Norway, employing ocean-bottom seismometers, Geophys. J. Int., Jones, C., Kanamori, H. & Roecker, S., Missing roots and mantle drips : regional Pn and teleseismic arrival times in the southern Sierra Nevada and vicinity, California, J. geophys. Res., 99(B3), Kaban, M.K.S., P., Artemieva, I.M. & Mooney, W.D., Density of the continental roots: compositional and thermal constraints, Earth planet. Sci. Lett., 209, Kanestrøm, R., Seismic investigations of the crust and upper mantle in Norway, in Deep Seismic Soundings in Northern Europe, pp , ed. Vogel, A., Swedish Natural Science Research Council, Stockholm. Kanestrøm, R. & Haugland, K., Profile section 3-4, in Deep Seismic Sounding in Northern Europe, pp , ed. Vogel, A., University of Uppsala. Kearey, P., Brooks, M. & Hill, I., An Introduction to Geophysical Exploration, 3rd edn, Blackwell Science, Oxford. Kinck, J.J., Husebye, E.S. & Larsson, F.R., The Moho depth distribution in Fennoscandia and the regional tectonic evolution from Archean to Permian times, Precambrian Res., 64, King, S.D. & Anderson, D.L., Edge-driven convection, Earth planet. Sci. Lett., 160, Koistinen, T., Stephens, M.B., Bogatchev, V., Nordgulen, Ø., Wennerstrøm, M. & Korhonen, J., Geological map of the Fennoscandian Shield, scale 1: , Geological Survey of Finland, E. G. S. o. N., Trondheim; Geological Survey of Sweden, Uppsala; Ministry of Natural Resources of Russia, Moscow. Koons, P.O., Three-dimensional critical wedges: tectonics and topography in oblique collisional orogens, J. geophys. Res., 99, Korja, A., Korja, T., Luosto, U. & Heikkinen, P., Seismic and geoelectric evidence for collisional and extensional events in the Fennoscandian Shield implications for Precambrian crustal evolution, Tectonophysics, 219, Journal compilation C 2009RAS

11 Korsman, K., Korja, T., Pajunen, M., Virransalo, P. & GGT/SVEKA Working Group, The GGT/SVEKA transect: structure and evolution of the continental crust in the Paleoproterozoic Svecofennian orogen in Finland, Int. Geol. Rev., 41, Lawrence, J.F., Wiens, D.A., Nyblade, A.A., Anandakrishnan, S., Shore, P.J. & Voigt, D., Crust and upper mantle structure of the Transantarctic Mountains and surrounding regions from receiver functions, surface waves, and gravity: implications for uplift models, Geochem., Geophys., Geosyst., 7, Lidberg, M., Johansson, J.M., Scherneck, H.-G. & Davis, J.L., An improved and extended GPS-derived 3D velocity field of the glacial isostatic adjustment (GIA) in Fennoscandia, J. Geodesy, 81, Lidmar-Bergstrøm, K., Ollier, C.D. & Sulebak, J.R., Landforms and uplift history of southern Norway, Global Planet. Change, 24, Ludwig, W.J., Nafe, J.E. & Drake, C.L., Seismic refraction, in The Sea, pp , ed. Maxwell, A.E., Wiley-Interscience, New York. Lyngsie, S.B., Thybo, H. & Lang, R., Rifting and lower crustal reflectivity: a case study of the intracratonic Dniepr-Donets rift zone, Ukraine, J. geophys. Res., 112, doi: /2006jb Milne, G.A., Mitrovica, J.X., Scherneck, H.-G., Davis, J.L., Johansson, J.M., Koivula, H. & Vermeer, M., Continuous GPS measurements of postglacial adjustment in Fennoscandia, 2. Modeling results, J. geophys. Res., 109, doi: Mykkeltveit, S., A seismic profile in southern Norway, Pure appl. Geophys., 118, Neumann, E.R., Olsen, K.H., Baldridge, W.S. & Sundvoll, B., The Oslo Rift: a review, Tectonophysics, 208, Nielsen, C. & Thybo, H., No Moho uplift below the Baikal Rift Zone: evidence from a seismic refraction profile across southern Lake Baikal, J. geophys. Res., in press, doi: /2008jb Niskanen, E., On the upheaval of land in Fennoscandia, Ann. Acad. Sci. Fenn Ser. A, 53, Olesen, O. et al., Bridging the gap between the onshore and offshore geology in Nordland, northern Norway, Nor. J. Geol., 82, Pascal, C. & Olesen, O., Are the Norwegian mountains compensated by a mantle thermal anomaly at depth?, Tectonophysics, in press, doi: /j.tecto Perez-Gussinye, M., Lowry, A.R., Watts, A.B. & Velicogna, I., On the recovery of effective elastic thickness using spectral methods: examples from synthetic data and from the Fennoscandian Shield, J. geophys. Res., 109, doi: Poudjom Djomani, Y.H., Fairhead, J.D. & Griffin, W.L., The flexural rigidity of Fennoscandia: reflection of the tectonothermal age of the lithospheric mantle, Earth planet. Sci. Lett., 174, Ramberg, I.B., Gravimetry interpretation of the Oslo Graben and associated igneous rocks, Nor. Geol. Unders., 325, Ramberg, I.B. & Smithson, S.B., Gravity interpretation of the Southern Oslo Graben and adjacent Precambrian rocks, Norway, Tectonophysics, 11, Riis, F., Quantification of Cainozoic vertical movements of Scandinavia by correlation of morphological surfaces with offshore data, Global Planet. Change, 12, Riis, F. & Fjeldskaar, W., On the magnitude of the Late Tertiary and Quaternary erosion and its significance for the uplift of Scandinavia and the Barents Sea, Structural and Tectonic Modelling and its Application to Petroleum Geology. NPF Special Publication, 1, Roberts, D. & Gee, D.G., An introduction to the structure of the Scandinavian Caledonide Orogen, John Wiley & Sons, Chichester. Moho map, southern Norway 1765 Sain, K. & Reddy, P.R., Use of post-critical reflections in solving the hidden layer problem of seismic refraction work, Geophysics, 62, Scherwath, M., Stern, T.A., Davey, F.J., Okaya, D., Holbrooke, W.S., Davies, R. & Kleffmann, S., Lithospheric structure across oblique continental collision in New Zealand from wide-angle P-wave modeling, J. geophys. Res., 108(B12), 2566, doi: /2002jb Sellevoll, M.A. & Warrick, R.E., A refraction study of the crustal structure in southern Norway, Bull. seism. Soc. Am., 61, Stephens, M.B., The Scandinavian Caledonides: a complexity of collisions, Geol. Today, 4, Stern, T.A., Gravity anomalies and crustal loading at and adjacent to the Alpine Fault, New Zealand, N. Z. J. Geol. Geophys., 38, Stern, T.A. & ten Brink, U.S., Flexural uplift of the Transantarctic Mountains, J. geophys. Res., 94, Stern, T.A., Baxter, A.K. & Barrett, P., Isostatic rebound due to glacial erosion within the Transantarctic Mountains, Geology, 33, Stewart, J. & Watts, A.B., Gravity anomalies and spatial variations of flexural rigidity at mountain ranges, J. geophys. Res., 102, Studinger, M., Bell, R.E., Buck, W.R., Karner, G.D. & Blankenship, D.D., Sub-ice geology inland of the Transantarctic Mountains in light of new aerogeophysical data, Earth planet. Sci. Lett., 220, Svenningsen, L., Balling, N., Jacobsen, B.H., Kind, R., Wylegalla, K. & Schweitzer, J., Crustal root beneath the highlands of southern Norway resolved by teleseismic receiver functions, Geophys. J. Int., 170, ten Brink, U.S., Hackney, R., Bannister, S., Stern, T.A. & Makovsky, Y., Uplift of the Transantarctic Mountains and the bedrock beneath the East Antarctic icecap, J. geophys. Res., 102, Thybo, H., Crustal structure along the EGT profile across the Tornquist Fan interpreted from seismic, gravity and magnetic data, Tectonophysics, 334, Thybo, H. & Nielsen, C.A., Magma-compensated crustal thinning in continental rift zones, Nature, 457, Thybo, H., Perchuc, E. & Zhou, S., 2000a. Intraplate earthquakes and a seismically defined lateral transition in the upper mantle, Geophys. Res. Lett., 27, Thybo, H., Maguire, P.K.H., Birt, C. & Perchuc, E., 2000b. Seismic reflectivity and magmatic underplating beneath the Kenya Rift, Geophys. Res. Lett., 27, Tryti, J. & Sellevoll, M.A., Seismic crustal study of the Oslo Rift, Pageoph, 115, Turcotte, D.L. & Schubert, G., Geodynamics: Applications of Continuum Physcis to Geological Problems, John Wiley & Sons, New York. Vestøl, O., Determination of postglacial land uplift in Fennoscandia from levelling, tide-gauges and continuous GPS stations using least squared collocation, J. Geod., 80, Vogel, A. & Lund, C.-E., Profile section 2 3, in Deep Seismic Soundings in Northern Europe, pp , ed. Vogel, A., Swedish Natural Science Research Council, Stockholm. Watts, A.B., Isostasy and Flexure of the Lithosphere, Cambridge University Press, Cambridge. Weidle, C. & Maupin, V., An upper-mantle S-wave velocity model for Northern Europe from Love and Rayleigh group velocities, Geophys. J. Int., 175, Zelt, C.A. & Smith, R.B., Seismic travel time inversion for 2-D crustal velocity structure, Geophys. J. Int., 108, Zhu, L. & Kanamori, H., Moho depth variation in southern California form teleseismic receiver functions, J. geophys. Res., 105, Ziegler, P.A., Evolution of the Arctic-North Atlantic and the Western Tethys, Am. Assoc. Petrol. Geol. Mem., 43,