New Moho map for onshore southern Norway

Geophys. J. Int. (2009) 178, 1755–1765 doi: 10.1111/j.1365-246X.2009.04240.x

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New Moho Map for onshore southern Norway

Wanda Stratford,1 Hans Thybo,1 Jan Inge Faleide,2 Odleiv Olesen3 and Ari Tryggvason4

1Department of Geography and Geology, University of Copenhagen, Copenhagen, Denmark. E-mail: [email protected] of Geosciences, University of Oslo, Oslo, Norway3Geological Survey of Norway, Trondheim, Norway4Department of Earth Sciences, Uppsala University, Uppsala, Sweden

Accepted 2009 May 6. Received 2009 May 3; in original form 2008 August 6

S U M M A R YA recent seismic refraction study across southern Norway has revealed that the up to 2469 mhigh 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. Evidencefrom new seismic data indicate that beneath the highest topography Moho depths are around38–40 km. These measurements are ∼2 km deeper than early estimates interpolated fromcoarsely spaced refraction profiles, but up to 3 km shallower than Receiver Function estimatesfor the area. Moho depth variation beneath the mountains roughly correlates with changes insurface topography indicating that topography is, at least to the first order, controlled by crustalthickness. However, the highest mountains do not overlie the thickest crust and additionalsupport for topography, for example from flexural strength in the lithosphere, low densitiesin the upper-mantle or mantle dynamics, is likely. The relationship between topography andMoho depth breaks down for the Oslo Graben and the Fennoscandian Shield to the eastand north. High density lower crustal rocks below Oslo Graben and increasing crust andlithospheric thicknesses below the Fennoscandian Shield may produce a negative correlationbetween topography and Moho depth.

Key words: Controlled source seismology; Continental margins: divergent; Dynamics oflithosphere and mantle; Crustal structure.

1 I N T RO D U C T I O N

Over the last 30 yr seismic studies of the lithosphere in southernNorway have led to models that predict a range of Moho depthsbeneath the Southern Scandes Mountains (Fig. 1b). Early stud-ies relied on coarse refraction profiling to sample Moho depths(Sellevoll & Warrick 1971; Kanestrøm 1971; Tryti & Sellevoll1977; Mykkeltveit 1980; Cassell et al. 1983) and interpolation be-tween surveys to infer regional trends (Sellevoll & Warrick 1971;Kinck et al. 1993; Fig. 1b). Large receiver spacing and gaps in thecoverage of these early studies meant that high resolution imag-ing of the Moho beneath the mountains was lacking. Results fromearly refraction profiling indicate that the crust is only 38 km thick(Sellevoll & Warrick 1971; Fig. 1b), and therefore, mechanismsother than variations in crustal thickness have been inferred forisostatic support of the observed topography (Olesen et al. 2002;Ebbing & Olesen 2005; Ebbing 2007).

A Bouguer gravity anomaly of –80 to –130 mGal, centred onthe Southern Scandes Mountains, is suggestive of a substantialdensity anomaly at depth. A correlation between topography andBouguer gravity has been used to infer Airy-isostatic support forthe mountains (Olesen et al. 2002; Ebbing & Olesen 2005) and thepresence of a crustal root (Balling 1980). Recent Receiver Func-tion studies (Fig. 1b) supported this argument as contrary to the

early refraction profiling, a crustal root and Moho depths of up to43 km are interpreted beneath the mountains (Svenningsen et al.2007).

A negative Bouguer gravity anomaly, however, does not nec-essarily demonstrate the presence of a crustal root, as the depthdistribution of densities is unknown and there can be dynamic sup-port from the mantle. Bouguer anomalies of similar magnitude andsign have been observed in other mountain ranges with ∼1200 mmean elevations such as the Southern Alps in New Zealand (Koons1994; Stern 1995) and the northern Transantarctic Mountains (tenBrink et al. 1997; Studinger et al. 2004; Stern et al. 2005). Despitethe similar mean elevations and Bouguer gravity signatures, theSouthern Alps have an over thickened crustal root and high man-tle densities to balance the surface topography (Scherwath et al.2003), whereas the Transantarctic Mountains are supported, at leastin part, by a low density mantle (Stern & ten Brink 1989; Lawrenceet al. 2006). For the Southern Scandes, flexural modelling (Ebbing& Olesen 2005) has been used to infer that the gravity anomaly isproduced, at least in part, by low densities below the Moho.

Presented here are new results from Magnus-Rex (Fig. 1a), arecent refraction seismic study across southern Norway. The studywas undertaken to add required high resolution constraints to, andupdate, the Moho depth map for the Southern Scandes Mountainsand to determine if the mountains are supported by a crustal root.

C© 2009 The Authors 1755Journal compilation C© 2009 RAS

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 multidisciplinarystudy of the lithosphere in southern Norway. Magnus-Rex seismic experiment: Red stars are shot locations, blue dots are seismograph stations. Line numbersand 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-Arsund (Mykkeltveit 1980).(E) Flora-Asnes (Sellevoll & Warrick 1971). (F) Oslo Graben (Tryti & Sellevoll 1977). (G) Trondheim-Sundsvall (Vogel & Lund 1971). (H) A crustal scalerefraction 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 theregions where upper allochthon rocks are still present. Blue shading outlines the Oslo Graben.

2 T E C T O N I C S E T T I N G

Bordered to the west by the passive margin of the North Atlantic andto the east by thick crust of the Fennoscandian Shield, the South-ern Scandes Mountains have undergone a number of uplift andsubsidence events (Lidmar-Bergstrøm et al. 2000; and referencestherein). Folding and faulting were induced by the collision betweenBaltica and Laurentia during the Caledonian orogeny from 440 to410 Ma (Andersen 1998). During this period of mountain building,allochthonous nappes were over-thrust onto Fennoscandian base-ment rock from the west (Roberts & Gee 1985). The nappes can bedivided into four main units, of which, three are present in southernNorway. Upper Allochthon Caledonides are preserved as far southas 62◦ N (Andersen 1998). A corridor of predominantly middle andlower allochthon Caledonides remain within the Faltungsgraben, adown-faulted graben associated with the postorogenic extensionalcollapse of the Caledonides in the southwest (Fig. 1b). The mid-dle and lower allochthons represent the nappe sequences formedby collision of the continental margin of Baltica. In contrast, theupper allochthon is comprised predominantly of sedimentary andigneous rocks derived from the pre-collision crust of the Iapetusocean, and include ophiolites and island arc complexes (Stephens1988). Outside the Caledonides, basement rock in southern Norwayis exposed and is predominantly comprised of granites deformed inthe Sveconorwegian orogeny (Koistinen et al. 2001).

Following the Caledonian orogeny, there was a reversal in defor-mation polarity and the Scandes underwent extension and postoro-genic collapse (Andersen 1998). Despite this extension, a long lived

mountain range throughout the Palaeozoic and Mesozoic has beenargued (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 incisionalerosion and other unknown mechanisms in the Neogene (Riis 1996;Lidmar-Bergstrøm et al. 2000; Faleide et al. 2002) and recent upliftof 1–4 mm yr−1 due to postglacial rebound from removal of theFennoscandian ice sheet (Niskanen 1939; Balling 1980). Althoughpostglacial rebound and incisional erosion are still in effect forthe present day surface dynamics (Riis & Fjeldskaar 1992), othertectonic processes may be required to explain the inferred longerterm uplift.

3 P R E V I O U S S T U D I E S

Refraction profiling began in earnest in the southern Norway regionin the early 1970s. These studies used small numbers of analogueinstruments that were relocated between shots to obtain the longoffsets required to record Pn arrivals (Sellevoll & Warrick 1971;Mykkeltveit 1980; Cassell et al. 1983). The velocity models de-rived for the crust from these early studies are generally of two orthree homogenous layers. Nevertheless, these early profiles showconsistency between the observed velocity structures with depththroughout the southern Norway region. From these studies, ve-locities for the upper crust in southern Norway are of 6.25 ± 0.1km 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, thereis good agreement between crustal velocity profiles below 5 km

C© 2009 The Authors, GJI, 178, 1755–1765

Journal compilation C© 2009 RAS

Moho map, southern Norway 1757

Figure 2. Profiles of 1-D velocity structure with depth extracted from pre-vious 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 aremarked with black dots. A 1-D profile extracted from line 1, this study, isalso shown (location is marked with a white dot in Fig. 1a).

depth. The interpretation of the Moho depth from these studies wasin part questioned due to the coarse resolution of these early mod-els. With large receiver spacings, these surveys were susceptible tothe hidden layer problem whereby arrivals from thin layers at depthare not observed as first arrivals and depths can be over or under-estimated (Kearey et al. 2002). Gaps between the surveys also leftregions where significant extrapolation was required to produce theoriginal maps of contoured crustal thickness (Sellevoll & Warrick1971; Kinck et al. 1993; Korsman et al. 1999).

4 DATA A N D R E S U LT S

Magnus-Rex, a new crustal scale seismic refraction experiment, wasconducted in October 2007. Three ∼400 km long seismic lines weredeployed across southern Norway (Fig. 1a). The elevated centralSouthern Scandes were targeted with two profiles in a large X-pattern. An additional profile to the south extends east–west acrossthe Southern Scandes, through the Oslo Graben and into Sweden(Fig. 1a). A total of 26 shots of 100–400 kg charge size were firedalong the three lines. End shots were 400 kg. Smaller charges of100 and 200 kg were distributed along the profiles. Ca. 750 verticalcomponent (Texan) seismographs were deployed at two km spacingalong the lines. A close instrument spacing of 750 m was used in a120 km section across the Oslo Graben. Processing of the seismic

data included basic frequency filtering (bandpass 2–4 to 12–25 Hz)and trace balancing.

The upper crust is well constrained by first arrival (Pg) refractionsfrom all shots. No distinct changes in slope on the Pg arrivals branchare observed (Fig. 3); rather there is a gradual increase in slope withoffset indicating an almost continuous increase in velocity withdepth. Apparent velocities are ∼5.8 km s−1 at near offsets gradingto 6.6 ± 0.1 km s−1 at offsets >150 km (Fig. 3). Velocities in themiddle crust are constrained by first arrival refractions beneath thecentre of the lines. The lower crust and the layer of high velocity(>7 km s−1) material at the base of the crust below the Oslo Grabenare constrained by the moveout of reversed PmP (Moho) reflections.Pn arrivals from beneath the centre of all three lines have apparentvelocities of 8.05 ± 0.1 km s−1 in all directions (Fig. 3) and crossoverdistances with Pg of 160–180 km.

The accuracy of picking seismic data depends on the noise leveland frequency of the arrivals and for crustal refraction data, arrivalscan 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, velocitiesare 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 & Smith1992) is undertaken to model the observed traveltimes of Pg, PmPand Pn. A 2-D velocity model is constructed from the observed trav-eltimes by projecting all arrivals onto a 2-D plane by conservation ofshot offset (Zelt & Smith 1992). Seismic models are constructed bya top down approach whereby near surface velocities and velocitystructures are modelled first.

Data-dependent errors are estimated from the root mean square(rms) misfit of the model to the traveltime picks, which is optimallywithin the pick uncertainties (Table 1). The normalized χ 2 or chi-squared statistic represents how well the model traveltimes fit theobserved traveltimes to within the pick uncertainties

χ 2 = 1

n − 1

n∑i=1

[(t(i)calc − t(i)obs)

σ (i)

]2

,

where t(i)calc and t(i)obs are the ith calculated and observed trav-eltimes, σ (i) is the traveltime uncertainty and n is the number ofobservations. Values less than 1 indicate an overfit of the model tothe data and X 2 values of 1 represent the best obtainable fit.

Model dependant uncertainties are estimated by forcing the raytracing solutions into end-member velocity models. The upper crustand middle crust contribute the least to the overall uncertainty inMoho depth as there is complete, reversed ray coverage from allshots. However, there is no first arrival coverage of the high-velocitylower crust. Thus, bounds on the depth to the Moho are estimated bythe degree to which the structure of this high-velocity layer can bealtered, while still retaining a fit to the observed traveltimes withinthe pick uncertainty (Table 1). Where there is coverage of Pg, PmPand Pn the uncertainty in Moho depth is ±1 km. Where constraintis from Pg and PmP only, the uncertainty is ±2 km.

In refraction profiling hidden layers can cause errors in depthdeterminations. However, the joint analysis of wide-angle reflec-tion and refraction data can limit the range of possible hidden layerthicknesses and velocities (Sain & Reddy 1997). For southern Nor-way, appropriate end-member velocities for possible low- and high-velocity layers are inferred to be 5.4 (e.g. metagreywacke) and7.4 km s−1 (gabbro; Christensen & Mooney 1995; Koistinen et al.2001). Applying these velocities, hidden high- and low-velocity lay-ers in the crust up to 3 km thickness produce Moho depths within the

C© 2009 The Authors, GJI, 178, 1755–1765



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: Crustalrefraction (Pg) P-waves velocities (Vp) of 5.8–6.6 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 Vp of ∼8.05 km s−1. (b) Shot 18, Line 2: Pg (5.8–6.4 km s−1), PmP and crustal reflectivity (PiP), Pn arrivals at offsets of∼180 km with apparent Vp of 8.05 km s−1. (c) Shot 26, Line 3: Pg (5.8–6.4 km s−1), PmP reflections, Pn arrivals at offsets of ∼180 km with apparent Vp of∼8.05 km s−1 and a strong reflection from the mantle at offsets of ∼300–450 km.

uncertainties bounds. Moreover, no traveltime skips can be seen inthe first arrival data that might elucidate a hidden layer of significantthickness.

5 C RU S TA L V E L O C I T Y S T RU C T U R E

Although crustal velocities beneath the Magnus-Rex lines appearto be almost laterally homogeneous, there are variations in thetop 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 top10 km is adequate to highlight significant lateral velocity varia-tions. Therefore, we attribute the observed velocity variations in thetop 5 km to the Caledonide nappe sequences on lines one and two,and on line three to the volcanic rocks and sediments within theOslo Graben.

Near surface velocities along line 1 decrease from ∼6 km s−1 inthe northwest to ∼5.8 km s−1 in the southeast (Fig. 4a). Althoughit is not possible to attribute these velocity changes to specific

C© 2009 The Authors, GJI, 178, 1755–1765



Table 1. Assigned pick uncertainties and calculated modelmisfits (rms and χ2 values): n is number of observations, σ

is assigned pick uncertainty, �trms is the rms misfit of themodel to the picks, and X 2 is the normalized chi-squaredvalue.

n σ (s) �trms (s) X 2

Line 1Pg 731 0. 1 0.117 1.37PmP 101 0.1 0.135 1.84Pn 36 0.1 0.061 0.38Line 2Pg 693 0.1 0.15 2.25PmP 223 0.1 0.176 3.11Pn 26 0.1 0.175 3.19Line 3Pg 845 0.1 0.139 1.93PmP 307 0.1 0.162 2.63Pn 112 0.1 0.15 2.27

composition changes, the profile line crosses out of Sveconorwe-gian basement rock (predominantly granite and granodiorite) in thenorthwest into the middle and lower allochthon of the Caledoniannappe sequences in the southeast.

Near surface velocities along line 2 decrease from ∼6 km s−1

in the south to ∼5.8 km s−1 in the north (Fig. 4b). The profileline tracks along the western edge of the nappe sequence in thesouthwest and near the crossover point with line 1, passes into theupper allochthon in the north. The high velocity body with Vp of∼6.5 km s−1 (at distances of 330–360 km on line 2) at a depth of3 km is possibly ophiolite rock associated with the upper allochthon(Roberts & Gee 1985; Andersen 1998). This high velocity body isobserved on shot 11 arrivals only, thus its lateral and depth extent arenot well constrained. Previous seismic surveying in this area (lineD on Fig. 1b) also found anomalous velocities at depth beneath thenappe sequences in this region. A 4 km thick low-velocity body at14 km depth has been inferred from an offset in first arrival refrac-tion on an array just to the southwest of shot 11 (Mykkeltveit 1980).Anomalous velocity-depth relationships may, thus, be a feature ofthis region of the Caledonides, although the different depth rangesof the putative bodies indicates that the variation may be attributedto more than one allochthon.

Near surface velocities on line 3 show the most variation in theOslo Graben (Fig. 4c). Velocities as low as 5.5 km s−1 here areattributed to Cambro-Silurian sediments and Permo-Carboniferousgranitic rocks within the graben (Neumann et al. 1992). In thewest near surface velocities are 5.8 km s−1 where the line crossesthe middle and lower allochthons. Between distances of 100 and240 km the near surface velocities are around ∼6.0 km s−1 withinthe Sveconorwegian basement.

Velocities in the upper to middle crust are more homogeneousalong 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. Thenotable exception is the Oslo Graben where higher velocities of6.6 km s−1 are inferred to occur at depths as shallow as 10 km.These high velocities are focused toward the western side of thegraben (Fig. 4c).

The lower crust is constrained by refracted arrivals on all threelines with velocities of 6.6 km s−1 at a depth of ∼22 km except forthe gap due to the Oslo Graben, where velocities at this depth are lesswell constrained (Fig. 4c). The lower-most crust where velocitiesof 7.0–7.1 km s−1 are inferred immediately above the Moho is notconstrained by refracted arrivals. Instead the thickness and velocity

of this layer has been determined from the best-fit model to explainthe moveout of reversed PmP reflections from the Moho. Variationsin Vp 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 ± 1 kmis modelled (Fig. 4a). A Moho depth of 39 ± 1 km is matchedon lines 1 and 2 at their cross point (Figs 4a and b). Constrainton Moho depth extends to the southwest of the crossing point onLine 2 and indicates a decrease in Moho depth to a flat section at38 ± 1 km at distances of 170–240 km and 36 ± 2 km by a distance of∼110 km from the south end of the line (Figs 4b). This result concurswith an earlier reflection profile along Sognefjord that led to theinterpretation of ∼31 km thick crust near the coast and 36 km thickcrust by 100 km inland (100 km from the west coast along profile H,Fig. 1b; Iwasaki et al. 1994). There is also consistency with resultsfrom two regional refraction profiles recorded in 1971 (Sellevoll &Warrick 1971) where depths of 38 km were inferred just north ofSognefjord (160 km distance along profile E, Fig. 1b) and 36 kmat ∼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 pointbetween the lines 1 and 2 (Fig. 4a).

As the Oslo Graben is relatively narrow, the profile length in thegraben is insufficient for first arrivals from the lower crust to beresolved. Thus, less constraint for lower crustal structure is avail-able from Line 3 and results from a previous north–south orientedseismic line (Tryti & Sellevoll 1977; profile F, Fig. 1b), are there-fore, included in the velocity modelling (Fig. 4c). Tryti & Sellevol(1977) inferred velocities of 7.1 km s−1 at just over 20 km depth inthe graben. Using this added constraint, the Moho up-warp beneaththe graben is interpreted to be small, of the order of ∼2 km, givinga Moho depth of 34 ± 2 km, which is in accord with the interpreta-tion of Tryti & Sellevol (1977). West of the Oslo Graben the Mohodepth is 36 ± 1 km at 180 km and 38 ± 1 km by 140 km from thewest end of the profile line (Fig. 4c).

Earlier studies of the Graben interpreted shallower Moho depthsof ∼30–32 from reflection (Kanestrøm & Haugland 1971) andgravity modelling (Ramberg & Smithson 1971). However, fromthe late 1970s there have been a number of studies that indicate asmaller 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 Mohotopography 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 ofmagma-compensated crustal thinning, whereby intruding magmacompensates for the crustal thinning taking place due to extensionand stretching.

6 N E W M O H O M A P F O R S O U T H E R NN O RWAY

Moho depth measurements from Pn and PmP arrivals are plotted attheir geographic locations and contour lines are hand drawn betweenthe values (Fig. 5). The simplest means of extrapolation are used andthe contouring process aims to limit extra structure. However, thereare still gap regions within the data set. Under regions where seismicconstraints from Pn arrivals from this study are applied, the Mohodepth map has uncertainties of ±1 km, where PmP arrivals onlyavailable the uncertainty is ±2 km. Where Moho depth constraints

C© 2009 The Authors, GJI, 178, 1755–1765



Oslo Graben

19 20 21 22 23 24 25 26

West East

7.1

6.6

6.3

6.0

6.6

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Velocities in km/s

8.05

1 2 3 4 5 6 7 8 9

101112131415161718 4(2)

6.8

6.4

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6.4

6.8

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Line 1

NorthSouth

7.1

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Velocities in km/s

6.2

6.5

8.05

5.8

6.2

6.2

6.15

Line 2

West East

7.1

6.6

6.3

6.0

6.3

Velocities in km/s

8.05

5.8

Velocities in km/s

Velocities in km/sLine 1

Line 2

Line 3

a.)

b.)

c.)

Moho

Moho

Moho

5.9

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 approachto fitting picked arrival times. Solid lines indicate where constraint is available from refractions. Thick dashed lines indicate where constraint is from PmPreflections 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. Thesurface 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 inFig. 1b), uncertainty in Moho depth is more difficult to quantifyas 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 PmPconstrained Moho depths in this study.

Overall, the new compilation map of Moho depths for onshoresouthern Norway is similar to that of Kinck et al. (1993), but the

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Figure 5. (a) New Moho map for southern Norway. Labelled contoursrepresent Moho depth in km. The map contains most of the original depthmeasurements used in the map of (Kinck et al. 1993; Fig. 1b), as well as newmeasurements from this study. Singular black dots on the map are locationsof depth measurements from early studies (see Fig. 1b for references andFig. 2 for velocity information from these lines). Lines of dots are theseismic profiles of Magnus-Rex. Black dots are the region of Pn Mohodepth constraint; Grey dots are where PmP depth constraints are available.The contour lines are superimposed on a map of topography that has beensmoothed by averaging elevations over a 50 km window. Note there is somecorrelation between topography and Moho depth beneath to the SouthernScandes; however the highest mountains are offset to the west of the thickestcrust.

extra constraints from this study have added details and improvedthe reliability significantly. The offshore data included in Kincket al. (1993) Moho map in Fig. 1(b), are not included in this plotas more recent studies using expanding spread profiles (ESP) ofthe Moho depth off the west coast provide better constraint onthe crustal structure here (Christiansson et al. 2000). From Chris-tiansson et al. (2000), seismic study a high velocity body in thelower crust near the coast was inferred and interpreted as eclogizedrock (Christiansson et al. 2000). Eclogites are found at the surfacein western Norway (Andersen 1998), however whether the puta-tive high velocity body of eclogite extends onshore into the lowercrust cannot be constrained by the new data presented here. Plannedonshore–offshore profiling may update the gradient in crustal thick-ness between onshore and the continental slope.

New constraints provided by this study cover some of the gapsbetween the earlier refraction profile studies in southern Norway.The original Moho map remains unchanged along the lines of theprior studies and where there is overlap between the prior studiesand the new data presented here, there is agreement on Moho depthwithin the uncertainties. The main difference is an extension of theslightly thicker crust (38 km) southward under the Southern ScandesMountains (compare the 38 km contour in Fig. 1b with Fig. 5).

A recent Receiver Function study along a profile line of similarcoverage to the Magnus-Rex Line 1 has been used to infer the pres-ence of a crustal root beneath the Southern Scandes (Svenningsenet 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 ReceiverFunction study concurs with that determined from the new activesource refraction results presented here; but the depth to the inter-face from refraction profiling is around 3 km less. The differencesin inferred Moho depth may, in part, be due to the assumed P-wavevelocity model for the lower crust used in the Receiver Functiondepth migration being 3 per cent higher than what has been foundfrom refraction profiling. This difference in velocity, however, canexplain only 200 m of the depth difference between the two tech-niques. The remaining difference is at the extreme bounds of thecombined uncertainties (±2 km) of the Receiver Function studyand (±1 km) of the refraction profiling.

Receiver Function techniques utilise the travel times of wavesconverted from P to S at the Moho and there is an inherent tradeoff between velocity and depth. Moreover, the trade off betweenvelocity and depth is particularly strong for VS compared to VP

and changes in Vp/VS of only 0.1 can lead to ∼4 km change incrustal thickness (Zhu & Kanamori, 2000). The earthquake wavesused in a Receiver Function study are also of significantly lowerfrequency, and hence lower resolution, than those frequencies usedin active source refraction studies. Differences in frequency contentbetween the two data sets and the influence of the thickness of theMoho boundary may be other factors that contribute to interpreta-tion differences. Therefore, we favour the shallower Moho depthsdetermined by refraction profiling in this study.

7 A I RY- I S O S TA S Y

We assess the degree of crustal support for topography by compar-ing the relationship between topography and crustal thickness to afirst-order model of crustal buoyancy (isostatic Moho, Fig. 6). Herecrustal thickness is taken as the depth to the refraction Moho. As-suming Airy-isostasy and complete compensation for topographyby the crust, total crustal thickness (ht) is related to topography (hc)by

ht = hi + hc

(1 + ρt

ρm − ρc

),

where hi is the crustal thickness at sea level for southern Norway(30 km; Kinck et al. 1993), ρc (crust) and ρt (topography) are theaverage density of the crust (2830 kg m−3, estimated from the av-erage crustal velocity from this study and an empirical relationshipbetween seismic velocity and density from Brocher (2005) usingthe original data of Ludwig et al. (1970)) and ρm is the density ofmantle lithosphere (3300 kg m−3). We use smoothed topography,with a 50 km diameter for the smoothing window, for hc. This for-mula assumes that the vertical density distribution in the crust is thesame everywhere along the profile.

An average crustal density is often used for isostatic calculationsand for southern Norway gives a density contrast at the Moho of470 kg m−3, which is larger than can be inferred from the lowercrustal and upper-mantle seismic velocities in this study. Moreover,if the compensation is assumed to be at the crust–mantle boundarythen the density contrast at this boundary will be the compensatingfactor. Thus, two models are used. One of average crustal den-sity (ρc and ρt = 2830 kg m−3) and one where topography (ρt =2670 kg m−3) is balanced by higher densities (ρc = 2950 kg m−3)

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30

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Figure 6. Graphs showing the Airy-Isostatic relationship between topography and Moho depth for southern Norway and south western Sweden. Moho depthsand topography are extracted from Fig. 5 along the seismic profiles. Moho depths from the map (Fig. 5) and predicted from the topography/Moho depthrelationship based on two Airy-isostatic balance models are shown. Isostatic Moho is the model using an average crustal density to model topography andIsostatic Moho (LC) refers to the model where a higher density lower crust (ρc = 2950 kg m−3) is used. Note the distinct differences in the relationshipbetween the Moho depth and smoothed topography for the Southern Scandes, and for the Oslo Graben (distances 225–275 km) and Swedish Fennoscandiancrust (distances 275–450 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 acrossthe Moho for southern Norway, that is in accord with the contrastinferred from seismic velocities near the Moho from this study.Both models are presented in Fig. 6 to illustrate the uncertainties inassuming densities with isostatic modelling.

Profiles comparing refraction Moho depth (from Fig. 5) andthe Moho depth predicted from the Airy-isostatic balancing ofsmoothed topography (eq. 1) are used to define the relationshipbetween topography and crustal thickness in southern Norway. 2-Dprofiles (Fig. 6) are extracted along the three Magnus-Rex seis-mic lines from the contours of the refraction Moho map and the

smoothed topography. This topography is smoothed with a 50 km di-ameter filter to remove short wavelengths; smoothing out the peaksand the deeply incised, but narrow, fiords. Wavelengths of ∼50 kmare significantly smaller than the topographic wavelengths that arelikely to be locally compensated.

There are uncertainties in the conversion of seismic velocities todensities and it is the pattern of the calculated isostatic compensationincluding form, wavelength and magnitude, which contains usefulinformation on how topography is supported.

Along line 1, the Airy-isostatic models have a maximum crustalthickness of ∼39 km, which is close to the 40 km interpreted for therefraction crustal thickness (Fig. 6a). However, the region of highest

C© 2009 The Authors, GJI, 178, 1755–1765



elevation is not located above the thickest crust, but rather is offset∼60 km to the west over a gradient in crustal thickness. An offsetbetween the highest elevation and the thickest crust is observed ata number of mountain ranges and has been attributed to the loadof the mountains being partially supported by the flexural strengthof the lithosphere (Forsyth 1985; Stewart & Watts 1997). East ofthe highest mountains the crustal thicknesses are higher than thoseinferred from Airy-isostatic balance of topography (Fig. 6a).

Line 2 is close to being in isostatic equilibrium for the southwestpart of the line, but again the highest elevations (around distancesof 180 km) remain under-compensated. The fit is poor at the north-eastern end of the line where the crust thickens to values higherthan the topographic elevations would suggest (Fig. 6b).

Along line 3 the deepest isostatic Moho is also offset to the westof the deepest refraction Moho beneath the mountains (Fig. 6c). Tothe east the low elevations are over-compensated by a thicker crustthan the isostatic Moho predicts. This is especially true for the OsloGraben where evidence from higher seismic velocities at shallowerdepths in the graben indicate that the crust here is of higher density,which has not been accounted for with the current calculations. Thelow topography of Sweden is also over-compensated with the den-sity structure of the isostasy models presented here. The indicatedover-compensation in the Fennoscandian Shield proper may also beexplained by the presence of a high density lower crust (BABELWorking Group 1993; Abramovitz et al. 1998) and a thicker litho-sphere (Thybo 2001; Kaban 2003; Artemieva 2007; Artemieva &Thybo 2008).

8 D I S C U S S I O N

Despite the high elevations of the Southern Scandes, the peaks arebounded by deeply incised valleys such that the mean elevations,when averaged over 50 km2, are only 1200–1500 m (Fig. 5). Theincised crust is slightly thinner than shield crust and relatively thinfor an orogenic belt (Christensen & Mooney 1995).

Simple 1-D isostatic calculations for southern Norway highlightthe difficulty in explaining topographic elevations in regions wherethere are lateral variations in crustal structure and elastic thickness.For southern Norway, the principal evidence for deviation fromAiry-isostasy (Fig. 6) is that the thickest crust is not located belowthe highest part of the southern Scandes Mountains.

Based on the assumed density model, the Oslo Graben and theFennoscandian Shield show no correlation with the models for Airy-isostatic balance (Fig. 6). It is likely that the crust beneath theOslo Graben may be of higher density than elsewhere in southernNorway (Ramberg 1976; Tryti & Sellevoll 1977; Ebbing et al.2005). Furthermore, crustal thicknesses are high (>40 km) in theSwedish part of the Fennoscandian Shield where topography haslow relief (<500 m; Korja et al. 1993; Fig. 5). A negative cor-relation between topography and Moho depth in the Oslo Grabenand Fennoscandian Shield is observed (Fig. 6). The Fennoscan-dian 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 effecton lithospheric buoyancy and, therefore, on topography.

For Airy-isostatic balance, however, the variation in thick-ness of the high velocity and hence high density, lowercrust in Fennoscandia needs to be considered (Korja et al.1993). A compilation map produced from seismic studies inFennoscandia has inferred that the high velocity (>7 km s−1) layerat the base of the crust varies between 0 and 25 km thick. Korja

et al.’s (1993) compilation map shows a general increase in thethickness of this layer to the north and east of the southern Scandes,with a thickness of 4–12 km beneath the mountains. However, onlylines C and B (Fig. 1b) east of the mountains in southern Norwaywere used in this compilation and the contours were extrapolated tothe west. A thinner high velocity lower crustal layer (<5 km thick)is inferred from this study and this is a significant contrast to thepredominantly 8–12 km thickness inferred for northern Norway andSweden (Korja et al. 1993). Where the high velocity (high density)lower crust is thicker a higher average crustal density should beused to isostatically balance topography.

Coherence techniques and gravity modelling have inferred ef-fective elastic thicknesses (Te) of 8–20 km for the Caledonides insouthern Norway (Poudjom Djomani et al. 1999; Perez-Gussinyeet al. 2004; Ebbing & Olesen 2005) increasing to ∼40 km in centralSweden (Perez-Gussinye et al. 2004). Te values beneath the south-ern Scandes are low to average for continental areas (Watts 2001).However, some degree of additional support from a lithosphere withelastic strength is inferred (Ebbing & Olesen 2005).

A combination of crustal thickness and some flexural supportmay, thus, explain a significant component of the topography insouthern Norway. However, better understanding of the elastic thick-ness variations across southern Norway is required before the mag-nitude of any additional topographic support mechanisms, such asadditional positive buoyancy from low mantle densities, can bequantified. Thus for the highest peaks of the southern ScandesMountains, full support of topography by crustal buoyancy cannotbe demonstrated and density differences within the upper mantlemay also be contributing (Olesen et al. 2002; Ebbing & Olesen2005; Ebbing 2007).

Further evidence for buoyancy within the upper mantle comesfrom earthquake Pn and Sn velocities. A zone of sub-Moho low ve-locities has been inferred to extend beneath the Southern Scandes(Bannister et al. 1991). Similar low-velocity mantle has been in-ferred beneath other mountain ranges where crustal thicknesses areinadequate to explain elevations such as the Transantarctic Moun-tains (Lawrence et al. 2006) and the southern Sierra Nevada (Joneset al. 1994).

Furthermore, effects on topography from dynamic processesmust be considered. Recent studies using GPS and precise lev-elling measurements have inferred uplift values of 0–3 mm per yrfor the southern Scandes (Danielsen 2001; Vestøl 2006; Lidberget al. 2006). Although these rates decrease towards the southwestin the southern Scandes, a continued contribution from the effectsof ice sheet removal may be in effect (Milne et al. 2004).

Other possible mantle sources for dynamic topography havebeen inferred from seismological studies (Weidle & Maupin 2008).These studies of mantle S and P-wave speeds have inferred a lowspeed zone in the upper mantle below depths of around 70 km(Weidle & Maupin 2008). This observation, combined with theearlier sub-Moho Pn and Sn study (Bannister et al. 1991), indi-cates that the mantle is anomalous beneath southern Norway. Suchlow mantle wave-speeds can be attributed to thermal or compo-sitional anomalies and, by applying thermal expansion (Turcotte& Schubert 1982), a low density mantle can be inferred. A re-cent study combining gravity and heat flow data of the Norwe-gian mountains, however, showed that there are no visible effectsof a mantle temperature anomaly at the surface (Pascal & Olesen2009). What is causing the inferred mantle wave speed anomalyand what effect it has on mantle buoyancy remains a critical ques-tion for the magnitude and timing of the uplift and of the southernScandes.

C© 2009 The Authors, GJI, 178, 1755–1765



New data presented here outline the crustal contribution to to-pography in southern Norway. A region of deep Moho (38–40 km)extends north–south beneath the Southern Scandes forming an ap-parent root between the extended crust of the Oslo Graben to theeast and the Atlantic Margin/North Sea to the west (Figs 1a and 5).The data indicate that the Moho is slightly deeper beneath the cen-tral high mountains, than elsewhere in southern Norway. A deeperMoho indicates that more isostatic support is possible from the crustthan previously inferred (Ebbing 2007). However, it is unclear howfar the relationship between topography and Moho depth might ex-tend. Flexural rigidity (Riis & Fjeldskaar 1992; Ebbing & Olesen2005) and the increasing lithospheric and crustal thickness to theeast and north beneath Fennoscandia (Calcagnile 1982) must betaken into account for assessments of isostatic balance of topogra-phy. Furthermore, the Southern Scandes are located at the transitionbetween the Fennoscandian Shield and the North Atlantic Ocean,which may cause dynamic forces from edge driven convection inthe mantle (King & Anderson 1998) or from release of compressivestresses associated with ridge push from the ocean into the shield(Thybo et al. 2000a). Regardless, the new Moho map presentedhere provides for the first time firm evidence on the relation be-tween crustal thickness and the present elevation of the SouthernScandes Mountains.

A C K N OW L E D G M E N T S

The instruments were provided by the PASSCAL facility of the In-corporated Research Institutions for Seismology (IRIS), PASSCALInstrument Centre, New Mexico Tech, Socorro and the Universityof Copenhagen. Data collected during this experiment will be avail-able through the IRIS Data Management Centre. Data acquisitionwas supported by the Danish National Science Research Council,the Carlsberg Foundation, the Norwegian Geological Survey andthe Norwegian Research Council. Appreciation is extended to theprincipal field technicians, G. Kaip, B. Greschke, P. Jørgensen, J.Gellein, L. Furuhaug, G. Storrø, T. Sørdal and A. K. Nilsen. Thanksalso to students from Copenhagen, Oslo, and Bergen universitiesand additional personal from the Norwegian Geological Survey whohelped with the field deployment. Appreciation is also extended toJ. Ebbing and R. England for constructive reviews of the paper.

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New Moho map for onshore southern Norway