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Precambrian Research 238 (2013) 186– 198

Contents lists available at ScienceDirect

Precambrian Research

journa l h om epa ge : www.elsev ier .com/ locate /precamres

Proterozoic boundary in southern Norway revealed byoint-inversion of P-receiver functions and surface waves

arianne L. Kolstrup ∗, Valérie Maupinepartment of Geosciences, University of Oslo, N-0371 Oslo, Norway

r t i c l e i n f o

rticle history:eceived 27 February 2013eceived in revised form5 September 2013ccepted 2 October 2013vailable online 12 October 2013

a b s t r a c t

The seismic S-wave velocity structure of the crust and uppermost mantle in southern Norway and westernSweden is estimated from joint inversion of teleseismic P-receiver functions and Rayleigh wave phasevelocities. The data comes from the temporary broad band seismological experiments MAGNUS andDANSEIS, and from permanent stations in Norway and Sweden. The S-wave velocity profiles in westernand central southern Norway show a sharp and well defined Moho discontinuity at depths between29 and 42 km, in contrast to the northeastern part of the study area where no sharp Moho is found.

eywords:ennoscandian Shieldouthwest Scandinavian-Domainrustal S-wave velocity structurerust-mantle-transition

The crust-mantle transition in this area is gradual and found at a depth from 40 to 50 km. The gradualcrust-mantle transition could be caused by mafic underplates that are only partially eclogitized, whereasthe sharp and well-defined Moho discontinuity can be a consequence of extensional collapse of theSveconorwegian orogen. The boundary between a sharp Moho discontinuity and a gradual crust-mantletransition has a northwest strike, which is approximately perpendicular to the Caledonides and suggests

ed th

oint inversion that the boundary surviv

. Introduction

Southern Norway is located on the western margin of theennoscandian Shield and records a geological history extend-ng from Paleoproterozoic times. Major events of crustal buildingnd reworking in the region are the magmatism associated withhe Transscandinavian Igneous Belt (TIB, 1.85–1.65 Ga), the accre-ion of the Gothian domain (1.75–1.5 Ga), and the Sveconorwegian1.25–0.9 Ga) and Caledonian (490–390 Ma) orogenies. The tec-onic setting of the Sveconorwegian orogeny is disputed (Slagstadt al., 2012), and the extent of transportation of terranes duringhis orogeny is unknown (Andersen, 2005; Bingen et al., 2008). Its not clear whether parts of southern Norway are exotic terranesBingen et al., 2008) or reworked Fennoscandian Shield (Andersent al., 2009b). Another question is the fate of the crust during sub-uction and exhumation in the Caledonian orogeny (e.g. Hackert al., 2010). A better knowledge of the structure and compositionf the crust and mantle can give important constraints on crustalormation and evolution.

Recent geophysical studies in the region have shown that thepper mantle below southern Norway has low seismic velocities

Medhus et al., 2009, 2012; Maupin, 2011; Wawerzinek et al.,013; Wawerzinek, 2012) that can be explained by a temperatureifference of about 200 ◦C (Goes et al., 2004) compared to the more

∗ Corresponding author. Tel.: +47 48032053.E-mail address: m.l.kolstrup@geo.uio.no (M.L. Kolstrup).

301-9268/$ – see front matter © 2013 Elsevier B.V. All rights reserved.ttp://dx.doi.org/10.1016/j.precamres.2013.10.004

e Caledonian orogeny.© 2013 Elsevier B.V. All rights reserved.

normal southern Swedish lithosphere. The crustal structure hasalso been studied by active seismic source refraction (Stratfordand Thybo, 2011b,a), surface wave tomography (Köhler et al.,2011, 2012) and P-receiver functions (Frassetto and Thybo, 2013),leading to three new Moho maps that agree with each other toa large degree. The resolution of the three methods are howeverquite different. While refraction seismics give precise informationalong a few lines, the surface wave tomography and the P-receiverfunctions provide a better coverage of the whole area. Surface wavetomography gives information on the average velocity with depth,while the P-receiver functions are well-suited to detect interfacesat depth but are dependent on a good knowledge of the averagevelocity to properly locate the depth of the interfaces. Surfacewaves and receiver functions are therefore two very complemen-tary data sets and their joint analysis has the advantage of providingaccurate depth profiles of the S-wave velocity over a given area.

In this study we investigate the seismic structure of the crustand uppermost mantle of southern Norway and western Swedenby jointly inverting the teleseismic P-receiver functions and theRayleigh wave dispersion curves obtained with data from the tem-porary broadband experiments MAGNUS (Weidle et al., 2010),DANSEIS, and from permanent stations in the area (Fig. 1). The jointinversion provides good resolution on the S-wave velocity structureof the lower crust and the crust-mantle transition in the region, and

the dense spacing of seismic stations gives the opportunity to studythe crust in its geological and tectonic context (Fig. 1). After presen-tation of the data, of the analysis procedure and of the geologicalsetting of the area, we will show how the profiles group according

M.L. Kolstrup, V. Maupin / Precambrian

Fig. 1. Left panel: Topographic map of the study area with location of seismologicalstations and the Norway-Sweden border. Stations are colour coded according tothe groups in Figs. 5, 6 and 10. Right panel: Outline of main geological units andboundaries in the Southwest Scandinavian Domain (SSD, Gaál and Gorbatschev,1987). Bamble Block (BB, nomenclature of Andersen (2005)), Caledonian Nappes(CN), Eastern Segment (ES, nomenclature of e.g. Bingen et al. (2005)), Hardanger-Rogaland Block (HRB, nomenclature of Andersen (2005)), Idefjorden Terrane (IT,nomenclature of e.g. Bingen et al. (2005)), Oslo Graben (OG), Telemark Block (TB,nomenclature of Andersen (2005)), Transscandinavian Igneous Belt (TIB), WesternGneiss Region (WGR). Mandal-Ustaoset Shear Zone (MUSZ), Mylonite Zone (MZ),Sveconorwegian Front (SF), Sveconorwegian Frontal Deformation Zone (SFDZ). (Forit

ta

2

Eistn

nterpretation of the references to colour in this figure legend, the reader is referredo the web version of this article.)

o geological regions and how this can give insight in the formationnd evolution of the area.

. Geological setting

The Fennoscandian Shield forms the northwestern part of theast European Craton. It contains an Archean core (3.7–2.6 Ga)

n the northeast and progressively younger domains towards theouthwest (Gaál and Gorbatschev, 1987). Island arcs that accretedo the Archean core in the Paleoproterozoic form the Svecofen-ian domain (1.91–1.87 Ga, Gaál and Gorbatschev, 1987; Nironen,

Research 238 (2013) 186– 198 187

1997). The Transscandinavian Igneous Belt (TIB) (1.85–1.65 Ga,Högdahl et al., 2004) of mainly alkali-rich granitoids extends fromsouthern Sweden to north-central Norway and separates the Sve-cofennian Domain to the east from the Southwest ScandinavianDomain to the west (SSD, Gaál and Gorbatschev, 1987) (Fig. 1).

The provinces of the Southwest Scandinavian Domain weredeformed and metamorphosed in the late Mesoproterozoic Sve-conorwegian orogeny (1.25–0.90 Ga, Gaál and Gorbatschev, 1987),but are mainly the products of earlier Mesoproterozoic continen-tal growth (Gaál and Gorbatschev, 1987; Bingen et al., 2005). TheSveconorwegian Frontal Deformation Zone (SFDZ, Wahlgren et al.,1994) constitutes the approximate eastern limit of Sveconorwe-gian reworking in the southeastern part of the study area, andthe Sveconorwegian Front (SF, Tucker et al., 1990) designates theapproximate limit of Sveconorwegian overprint in the northwest-ern part of the study area. The continuations of the SFDZ and theSF are hidden below Caledonian thrust sheets (Fig. 1). The MyloniteZone (MZ) constitutes a major Sveconorwegian shear zone and ter-rane boundary, separating the Paleoproterozoic TIB rocks of theEastern Segment (ES, Fig. 1) from the Gothian (1.75–1.5 Ga, Gaáland Gorbatschev, 1987) Idefjorden Terrane (IT, Fig. 1) to the west(Andersson et al., 2002).

During the Caledonian orogeny (490–390 Ma) crustal sheets andnappes were thrust over the western margin of the FennoscandianShield (Roberts and Gee, 1985). The Fennoscandian basement andparts of the allochtoneous nappes were subducted to great depthsand heavily deformed (Hacker et al., 2010, and references within).The transition between the Sveconorwegian provinces of south-ern Norway and the Caledonian Western Gneiss Region (WGR) isnow hidden below Caledonian allochtons. Likewise, the transitionbetween the Paleoproterozoic TIB rocks of the Eastern Segmentand the northern part of the WGR is covered by Caledonian nappes(Fig. 1).

The Oslo Graben formed during the late Carboniferous and earlyPermian (305–241 Ma, Neumann et al., 2004) and separates thenorthwest and southeast part of the Mesoproterozoic Idefjordenterrane (Fig. 1).

3. Data and methodology

In this work, we use seismological data recorded by the tempo-rary MAGNUS network between September 2006 and June 2008, bythe temporary DANSEIS network from April 2008 to June 2009, andby permanent stations in the study area (see Fig. 1). For some of thepermanent stations (KONO, NORSAR array, HFC2) data is availablefor more than 10 years.

We derive the Sv-wave velocity profile beneath each station ofthe network by combining the information obtained by P-wavereceiver functions and Rayleigh wave phase velocities in the periodrange 3 to 67 s. The principle of the P-wave receiver function is toextract from the coda of a teleseismic P-wave the arrival time andamplitude of S-waves and multiples converted from the main P-wave at interfaces below the station (Langston, 1979). The delaybetween the main P-wave and the converted waves depends on thedepth to the interface and the Sv-wave velocity above, whereas theamplitude of the converted waves is sensitive to contrasts in veloc-ities and density at the interface. Radial P-wave receiver functionsare estimated by deconvolving the vertical seismogram from thecorresponding radial seismogram (Langston, 1979; Ammon, 1991).The receiver function technique has been used extensively to studythe crust and Moho discontinuity (e.g. Langston, 1979; Owens et al.,

1984; Ammon et al., 1990; Ammon, 1991).

Rayleigh waves are surface waves that travel with phase veloc-ities that depend mostly on averages of the Sv-wave velocitiesin different depth ranges. Since receiver functions are sensitive

188 M.L. Kolstrup, V. Maupin / Precambrian Research 238 (2013) 186– 198

NWG

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Fig. 2. Accepted events at the temporary station

o discontinuities in the velocity profile and the Rayleigh waveelocities are sensitive to their average, a joint inversion of the twoata sets reduces the non-uniqueness of the inversions of each dataet taken separately and has become widely used (Özalaybey et al.,997; Du and Foulger, 1999; Julià et al., 2000). We use the joint

nversion procedure developed by Julià et al. (2000, 2003) in anpen source code by Herrmann and Ammon (2002) to jointly invertayleigh wave phase velocities and P-wave receiver functions.

The inversion procedure (Julià et al., 2000; Herrmann andmmon, 2002) assumes that the structure below the receiver can,

n an area of typically 40 × 40 km2 for Moho studies, be mod-lled by spherical (flat), isotropic layers. In case of dipping layersnd/or anisotropy, it has been shown that radial and transverseeceiver functions have a distinctive azimuthal variation (Cassidy,992; Savage, 1998; Frederiksen and Bostock, 2000). Hence, thezimuthal variation of the receiver functions is evaluated beforenversion at each station to determine if the assumption of a spher-cal, isotropic Earth below the receiver is reasonably satisfied.

.1. P-wave receiver functions

Receiver functions used in this study were calculated byrassetto and Thybo (2010, 2013) from teleseismic P and PPaveforms in a receiver function migration study using iterative,

ime-domain deconvolution (Ligorría and Ammon, 1999). Eventsrom great circle path distances of 25–95◦ with M > 5.6 for P arrivals,nd from distances of 95–180◦ with M > 6.4 for PP arrivals wereonsidered (Frassetto and Thybo, 2013).

Receiver functions with Gaussian width factors of 1 and 2.5 aresed here, giving sensitivity to structures thicker than approxi-ately 2.0 and 0.8 km for the crust and upper mantle (Frassetto

t al., 2011). We reject receiver functions that explain less than0% of the observed radial seismogram, and only select receiverunctions where the direct P arrival has a higher amplitude thanll subsequent arrivals, negative as well as positive, and is not pre-eded by a trough or delayed significantly (Frassetto et al., 2011).

For all temporary deployments we have good quality eventsrom the northeast quadrant in the distance range of 60–95◦, and at

any stations we also have data from the other quadrants. Two typ-

cal examples of event distribution with back azimuth are shownn Fig. 2 for a temporary station and a permanent one.

To improve the signal to noise ratio, receiver functions weretacked according to natural clusters in the data and in ray

16 (left) and the permanent station HFC2 (right).

parameter bins smaller than 0.01 s/km. For stations with a goodback azimuthal coverage, stacks were performed in small backazimuth bins (3–10◦) in order to evaluate the variation of bothradial and transverse receiver functions with back azimuth. Stackedreceiver functions were kept for inversion if they contained 3 ormore coherent traces. Stacks with two traces were kept only ifthe traces were extremely coherent. Table 1 lists the number ofreceiver function stacks used in the inversion for each station andthe corresponding range of ray parameters.

3.2. Surface waves

We use Rayleigh wave phase velocities from the surface wavetomographic study of Köhler et al. (2012). Phase velocities in theperiod range 3–30 s are derived from ambient noise data and in theperiod range 25–67 s from earthquake data (Köhler et al., 2012).Uncertainties lie in the range of 0.003–0.03 km/s for short periods,and in the range of 0.005–0.03 km/s for the longer periods. In thecentral areas of the study area, phase velocity maps are available atperiods from 3 to 67 s, giving information about velocities down todepths of 90–130 km. At the outskirts, periods are available from 3to 30 s, providing velocity information down to depths of 40–60 km.The range of periods available at each station is listed in Table 1.

Since the P-wave receiver functions are not sensitive to the Sh-wave velocity structure, we do not use in the present analysis theLove wave phase velocities measured for the study area in Köhleret al. (2012).

3.3. Joint inversion

The joint inversion of the P-wave receiver functions andRayleigh wave phase velocities is performed using an iterative,damped least-squares scheme (Julià et al., 2000; Herrmann andAmmon, 2002). The relative weighting of the two data sets in ajoint inversion at a given station is determined by the parameterp (Julià et al., 2000; Herrmann and Ammon, 2002). This parametertakes values between 0 and 1, where 0 corresponds to an inversionwith only receiver functions and 1 corresponds to an inversion withonly phase velocities. A damping factor determines the amount of

differential smoothing applied (Herrmann and Ammon, 2002), anda modest damping factor of 0.5 is used in this study. The startingmodel consists of 82 layers down to a depth of 570 km and fol-lows the AK135-F continental model (Kennett et al., 1995) below

M.L. Kolstrup, V. Maupin / Precambrian Research 238 (2013) 186– 198 189

Table 1Stations, data, parameters.

Station Number of invertedreceiver function stacks

Back azimuthalcoverage of stacks

Ray parameterrange (s/km)

Phase velocityperiod range

Weightingparameter p

Maximum velocitygradient (s−1)

Moho depth(km)

BER 10 N, NE, E, W 0.042–0.063 3–30 0.35 0.1436 30FOO 10 N, NE, E, S, SW, W 0.042–0.080 3–30 0.35 0.1113 33HFC2 25 N, NE, E, SE, SW, W 0.044–0.074 3–30 0.35 0.0558 NAHT2S 2 N, NE 0.055–0.058 3–30 0.5 0.0818 NAHT3S 3 N, NE, E 0.052–0.057 3–67 0.5 0.0828 NAHT4S 2 NE 0.054–0.057 3–67 0.5 0.0631 NAHT6S 4 N, NE, E, S 0.052–0.081 3–30 0.5 0.0686 NAKONO 28 N, NE, E, S, SW, W, NW 0.041–0.080 3–67 NA NA NAMOL 11 N, NE, E, S 0.043–0.081 3–67 0.35 0.0927 39NAO01 28 N, NE, E, S, SW, W, NW 0.043–0.081 3–67 0.35 0.1219 34NC204 18 N, NE, E, S, SW, W 0.044–0.081 3–67 0.35 0.0895 39NC405 19 N, NE, E, SW, W 0.044–0.081 3–67 0.35 0.0621 NANC602 23 N, NE, E, SE, S, SW, W,

NW0.043–0.081 3–67 0.35 0.0981 38

NWG01 5 N, NE, E, SW 0.044–0.058 3–67 0.5 0.0525 NANWG03 8 N, NE, E, SW, W 0.042–0.061 3–30 0.5 0.0501 NANWG04 2 N, NE 0.054–0.058 3–67 0.5 0.0678 NANWG05 4 N, NE, E 0.045–0.058 3–67 0.5 0.0566 NANWG06 4 NE, SW 0.046–0.058 3–67 0.5 0.0520 NANWG07 5 N, NE 0.047–0.060 3–30 0.5 0.0978 38NWG08 5 N, NE, E 0.045–0.066 3–67 0.5 0.0640 NANWG09 2 NE, SW 0.054–0.060 3–67 0.5 0.1389 40NWG10 5 N, NE, E 0.043–0.058 3–67 0.5 0.0457 NANWG11 2 NE, E 0.045–0.056 3–67 0.5 0.0737 NANWG12 3 N, NE, E 0.045–0.057 3–67 0.5 0.0905 40NWG13 4 N, NE 0.053–0.060 3–67 0.5 0.1353 38NWG14 5 NE, E, S 0.042–0.081 3–30 0.35 0.1205 31NWG15 4 N, NE, E, S 0.043–0.081 3–67 NA NA NANWG16 7 N, NE, E, S 0.043–0.082 3–67 0.35 0.1429 42NWG17/33 7 N, NE, SW 0.045–0.059 3–67 0.5 0.0996 40NWG18 8 N, NE, E, SW 0.045–0.067 3–67 0.35 0.0439 NANWG19 7 N, NE, E, S, W 0.045–0.059 3–67 0.5 0.1437 34NWG20 4 N, NE, E, SW 0.044–0.059 3–67 0.5 0.0788 20NWG21 4 N, NE, W 0.045–0.056 3–40 0.5 0.1083 36NWG22 4 N, NE, SW, W 0.045–0.061 3–30 0.5 0.1016 30NWG23/35 5 N, NE, E, W 0.047–0.067 3–30 0.5 0.0991 36NWG24 5 N, NE, E, SW 0.043–0.059 3–30 0.5 0.1498 36NWG25 3 N, NE, E 0.045–0.055 3–30 0.5 0.1088 34NWG26/34 7 N, NE, E, SW 0.043–0.060 3–50 0.5 0.1395 36NWG27 6 NE, E, SW 0.044–0.059 3–30 0.5 0.1102 34NWG28 6 N, NE, E, W 0.042–0.058 3–30 0.5 0.1453 32NWG29 5 N, NE, E, SW, W 0.043–0.060 3–30 0.5 0.0790 NANWG30 4 N, NE, E 0.042–0.055 3–30 0.5 0.1541 28NWG31 6 N, NE, E, SW, W 0.044–0.060 3–30 0.5 0.1094 34NWG32 7 N, NE, E, SW 0.043–0.060 3–67 NA NA NA

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ODD1 5 N, NE, E, W 0.047–TBLU 8 N, NE, E 0.044–

0 km. In the upper 50 km the velocities are simply replaced by theelocity value of AK135 at a depth of 50 km as in Herrmann andmmon (2002) and Yoo et al. (2007). Layer thicknesses are 2 km in

he upper 50 km, 5 km from 50 to 100 km depth, and 10 km below00 km depth. In the upper 50 km the model is allowed to changereely in the inversion, and between 50 and 90 km a tapered weight-ng is used to gradually disallow small-scale variations. Between 90nd 400 km only very slight changes are allowed.

An example of data and resulting model predictions at theemporary station NWG16 is displayed in Fig. 3. At this station weave phase velocities from 3 to 67 s period and 7 receiver functiontacks with a range of ray parameters from 0.04 to 0.08 km/sTable 1). The importance of the relative weighting of the two dataets is illustrated in Fig. 4, which shows inversions performed using-values varying from 0.05 to 0.95. Low p-values, correspondingo inversion based mostly on receiver functions, provide moremall-scale variations in the inverted model, but the cost is increas-

ng non-uniqueness as receiver function inversions inherentlyrovide non-unique models of velocity with depth (Ammon et al.,990), unless a specifically adapted model parametrization is usedJacobsen and Svenningsen, 2008). High p-values, corresponding

3–40 0.5 0.0945 36 3–30 0.35 0.0446 NA

to inversions based mostly on phase velocities, give a very smoothmodel that is not able to resolve sharp velocity changes. From theresolution kernels for the joint data set (not shown) we evaluatethat p-values between 0.35 and 0.5 provide a good compromisebetween uniqueness and detail in the inverted models. Evaluationof the resolution kernels also shows that for central stations withlong period surface waves, the resolution is good down to depthsof 60–65 km and for stations where the longest period is 30 s, theresolution is good down to 50–55 km.

The amount of receiver function stacks and the range of rayparameters available vary between stations (Table 1). When onlya few stacked receiver functions are inverted, it is more likely thatwe fit the model to local noise from one particular back azimuthand distance. Hence, the inverted models from stations with fewerreceiver functions appear slightly less smoothed. To ease the com-parison between models from stations with few and many receiverfunction stacks, we therefore increase the weight on the surface

wave data to 0.5 for stations with few receiver function stacks.This ensures a slightly smoother velocity model without applyingstronger differential damping in the inversion. The p-values appliedin each inversion can be found in Table 1.

190 M.L. Kolstrup, V. Maupin / Precambrian Research 238 (2013) 186– 198

F CentraR or intet

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ig. 3. Station NWG16. Left panel: Inverted model (red) and initial model (blue).

ayleigh wave phase velocities (black circles) and model predictions (red curve). (Fhe web version of this article.)

Uncertainty in the depth estimate to an interface is limitedy the model parametrization. With layers of 2 km thickness, thencertainty in depth is at best ±1 km in the upper 50 km of theodel and ±2.5 km below. Uncertainties in inverted velocities areore difficult to estimate, as we believe that the greatest uncer-

ainty does not come from the chosen inversion parameters, butrom the basic assumptions of a flat, isotropic Earth. At manytations, the back azimuthal coverage is sufficient to allow us tovaluate whether or not the data is suitable for inversion, but whereata is scarce there is no guarantee that the structure below thetation is simple enough for the formal inversion. Julià et al. (2005)stimated the velocity uncertainty to 0.1 km/s for the crust by using

range of inversion parameters and constraints. We regard this as minimum estimate in our study area.

.4. Characterization of the Moho discontinuity

Our initial model has no Moho discontinuity. Depending on theature of the data, and especially on the nature of the receiver func-ions, a clear Moho discontinuity may or may not develop in ournal models. In order to quantify the character of the crust-mantleransition, we calculate the maximum S-wave velocity gradient in

he lower crust and uppermost mantle. This maximum velocityradient is found to have values between 0.04 and 0.16 s−1 in thetudy area (Fig. 8). We define the Moho depth as the depth wherehe velocity gradient is maximum but restrict it to cases where

l panel: Receiver function stacks (blue) and model predictions (red). Right panel:rpretation of the references to colour in this figure legend, the reader is referred to

the maximum gradient is stronger than 0.085 s−1 and found in thedepth range where velocities are greater than 4.1 km/s. Consid-ering the velocities of 4.0–4.05 km/s found in the lower crust inthe refraction experiment of Stratford and Thybo (2011b), thisensures that the Moho is defined as a discontinuity if seismicvelocities increase from lower crustal ones to mantle ones in adepth interval not larger than about 5 km. To guide the eye, a linewith a constant velocity of 4.1 km/s has been added to all veloc-ity models and the velocities above this value shaded accordingto their values (e.g. Fig. 5). Where the maximum velocity gradientis smaller than 0.085 s−1 we define the crust-mantle transition asgradual.

4. Crustal structure from joint inversion

Out of 45 stations, 3 stations (NWG15, NWG32, KONO, Fig. 1,Table 1) exhibit strong systematic variation of the receiver func-tions with back azimuth, suggesting a complicated structure belowthe station, for example dipping layers, anisotropy, or structuralheterogeneity, that does not comply with our assumption of spher-ical and isotropic layers. We did therefore not perform a formalinversion for these stations. For the other 42 stations, receiver func-

tions exhibit minor variation with back azimuth, implying that thestructure below the station can be reasonably modelled using theinversion scheme of Julià et al. (2000) and Herrmann and Ammon(2002).

M.L. Kolstrup, V. Maupin / Precambrian

Fig. 4. Inverted models using p-values ranging from 0.05 to 0.95 (station NWG16).A p-value of 0 corresponds to an inversion with only receiver functions and 1cro

gpaWa

4

4

tn

orresponds to an inversion with only phase velocities. (For interpretation of theeferences to colour in this figure legend, the reader is referred to the web versionf this article.)

The 42 velocity profiles resulting from the inversion have beenrouped according to large scale geological units. In Fig. 6 weresent a summary with a profile from each unit, and in Figs. 5nd 10 we present all profiles from the Eastern Segment and theestern Gneiss Region, respectively. All inverted models, grouped

ccording to geological units, can be found in Appendix A.

.1. Northeastern Southern Norway

.1.1. Eastern Segment (TIB)Fig. 5 shows inverted models from the Eastern Segment (black

riangles in Fig. 1). The permanent station NC602 is located in theorthwestern part of the Eastern Segment, close to the Oslo Rift,

Fig. 5. Velocity profiles from stations in the Eastern Segment. Locations are de

Research 238 (2013) 186– 198 191

and the other permanent station HFC2 is located to the southeast.Temporary stations in this geological unit are NWG18, NWG11,HT3S and HT4S. The velocity profile from station NWG18 is shownin Fig. 6 plot A for comparison with profiles from other geologi-cal units. Details about the data sets available at each station andtheir relative weighting in the inversion can be found in Table 1.The velocity profiles for stations NWG18, NWG11, HT3S, HT4S andHFC2 are very similar. In the upper crust velocities increase fromapproximately 3.35 km/s at the surface to 3.6 km/s at 8 km depth.From 8 to 20 km velocities are approximately constant at 3.6 km/s.Between 20 and 30 km depth velocities increase again to about3.85 km/s and stay fairly constant in the depth range of 30–40 km.A gradual increase in velocities begins at depths greater 40 km andcontinues to 50 km, where the gradient in velocities decreases toapproximately zero. Typical mantle S-wave velocities greater than4.4 km/s are reached at depths around 50 km.

The velocity models show no clear discontinuity that could beidentified as the Moho. Instead, the profiles have a gradual tran-sition from crustal to mantle velocities from 40 to 50 km depth.The lack of a sharp Moho discontinuity in the profile results fromthe fact that the receiver functions contain only broad, low ampli-tude Ps conversions, as can be seen in the receiver functionsfrom station HFC2 in Fig. 7. The receiver functions have three lowamplitude Ps conversions at times of 2.0 s, 3.5 s and 6 s after thedirect P arrival. These conversions correspond to the three depthranges with positive gradients in velocity at 0–8 km, 16–24 km and44–50 km.

The lack of a clear Moho Ps conversion at station HFC2 was alsonoted by Svenningsen et al. (2007) in a receiver function migra-tion study along two lines in southern Norway. Svenningsen et al.(2007) had only two stations in the Eastern Segment and there-fore suggested that the lack of a clear Moho Ps conversion at HFC2was due to scattering of the Ps phase in the Mylonite Zone (MZ,Fig. 1), but also noted that it could be due to a gradual crust-mantletransition in the area.

The similarity between the receiver functions as well as theinverted models from station NWG18, NWG11, HT3S, HT4S andHFC2 suggests that the broad, low amplitude Moho Ps conversionsin the receiver functions is a stable feature, not related to noise orscattering, and that the gradual crust-mantle transition is a struc-

tural characteristic of this geological region. Another support forthis interpretation is the absence of a clear Moho Ps conversion fromany back azimuth at the permanent station HFC2 (Fig. 7) despite thegood data quality at this station.

noted by black triangles in Fig. 1. Velocities above 4.1 km/s are shaded.

192 M.L. Kolstrup, V. Maupin / Precambrian Research 238 (2013) 186– 198

F e Bloc( iss Re

arssiwl

ig. 6. Velocity profiles from representative stations in all geological units. BamblHRB), Idefjorden Terrane (IT), Oslo Graben (OG), Telemark Block (TB), Western Gne

The velocity profile from the westernmost station NC602 has shape significantly different from the other profiles (Fig. 5). Theeceiver functions at NC602 show a more distinct Moho Ps conver-ion at 4 s, which gives a sharper crust-mantle transition. The profile

hows a change in velocity gradient at depths of 36–38 km with anncrease in velocities from 4.0 to 4.35 km/s that can be associated

ith the Moho discontinuity. Additionally, there is a high velocityayer in the upper crust at depths of 12–18 km, which is not present

Fig. 7. Data from station HFC2. Left panel: stacked receiver functions ac

k (BB), Caledonian Nappes (CN), Eastern Segment (ES), Hardanger-Rogaland Blockgion (WGR).

in the other models in Fig. 5. The maximum S-wave velocity gradi-ent in the lower crust and uppermost mantle is 0.10 s−1 for stationNC602 at a depth of 38 km. For stations NWG18, NWG11, HT3S,HT4S and HFC2 the maximum velocity gradient has lower values,

between 0.044 and 0.083 s−1 (Table 1). The crust-mantle transitionis therefore defined as gradual at these stations. Thus, although thesurface geology places station NC602 along with the other stationsin the Eastern Segment, the S-wave velocity structure of the crust

cepted for inversion. Right panel: Rayleigh wave phase velocities.

brian Research 238 (2013) 186– 198 193

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stNalcNiilgatd

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Fig. 8. Maximum S-wave velocity gradient in the lower crust and uppermost man-tle. Black crosses show the location of stations where the joint inversion was notperformed. Black lines depict main geological units and boundaries from Fig. 1. Thesize of the circles represents approximately the size of the areas sampled by the

M.L. Kolstrup, V. Maupin / Precam

nd shallow upper mantle shows more similarities with the struc-ure below stations in the nearby Oslo Graben and in southwestorway (Section 4.2.2).

.1.2. Northern Western Gneiss Region and Caledonian NappesTwo other sets of inverted models show the same structural

haracteristics as the profiles from the the Eastern Segment. Thesetations are in the northeast part of the study area and cut throughhe Caledonian Nappes and the northern part of the Western Gneissegion (yellow and light blue triangles in Fig. 1). This part of theestern Gneiss Region comprises Caledonian reworked Paleopro-

erozoic basement (Tucker et al., 1990; Austrheim et al., 2003;ingen and Solli, 2009), but most of the area is covered by Cale-onian nappes and its history is therefore not well known.

Velocity profiles from stations NWG06 and NWG04 are pre-ented in Fig. 6 plots B and C, and models at all stations in thesewo areas can be found in Figures A.1 and A.2 in Appendix A. StationWG06 is located in the far northeastern part of our study area in

region covered by Caledonian nappes, while station NWG04 isocated closer to the Sveconorwegian Front (Fig. 1). We see from aomparison of plots A, B and C (Fig. 6) that the models from stationsWG06 and NWG04 are very similar to the model at station NWG18

n the Eastern Segment and have similar small velocity gradientsn the crust-mantle transition. Velocities are fairly constant in theower crust at depths of 30–40 km. Below 40 km, velocities increaseradually until typical mantle velocities are reached at depths ofround 46–50 km. For station NWG03, which is located at the coast,he gradual transition to mantle velocities happens at a shallowerepth from 30 to 40 km.

Station NC405 is located very close to the uncovered TIB rocksf the Eastern Segment and has a velocity profile very simi-ar in shape to the models in Fig. 5. Another interesting features that for stations NWG08 and NWG10 there is a tendency ofeeing a clearer (higher amplitude and smaller width) Moho Psonversion from data with back azimuth in the W and SW direc-ions, respectively, suggesting that these stations are close to theestern/southwestern edge of the area with gradual crust-mantle

ransition. Svenningsen et al. (2007) found a similar azimuthalariation at their station JH13, which was situated in the trian-le between NWG08, NWG10 and NWG12, and interpreted it as

Moho dipping northeastwards by about 10 degrees.Together with the stations of the Eastern Segment, the stations

n these two sets define a geographical area in the northeast-rn part of our study region where the crust-mantle transitions consistently gradual. This is clearly seen in Fig. 8 where grad-al crust-mantle transitions with vertical gradients lower than.085 s−1 are coloured light to dark blue whereas sharp Moho dis-ontinuities with gradients larger than 0.085 s−1 are shown withhite to red colours. In the Moho depth map in Fig. 9, stations with

radual crust-mantle transitions are marked by grey triangles.

.2. Southwestern Southern Norway

.2.1. Idefjorden terraneThe Idefjorden terrane (white triangles in Fig. 1) is split into

wo regions due to the crossing NNE-striking Oslo Graben. StationWG17/33 is located northwest of the Oslo Graben and stationsWG19 and NWG26/34 are located southeast of the Oslo Graben.eceiver functions from these three stations are very similar for allack azimuths and the data is well suited for inversion. All profilesre very similar (Figure A.3).

When we compare the model from station NWG19 in the Ide-

jorden terrane (Fig. 6 plot D) to models from the neighbouringastern Segment (Figs. 5 and 6 plot A) the striking difference ishe sharpness of the crust-mantle transition. The velocity profilesrom the Idefjorden terrane have a lower crust with a small velocity

receiver functions and the lateral resolution of the surface wave tomography. (Forinterpretation of the references to colour in this figure legend, the reader is referredto the web version of this article.)

gradient and a crust-mantle transition with a large velocity gradi-ent. For these three stations, velocities change from approximately3.95–4.40 km/s over a depth range of only 2–4 km with maxi-mum velocity gradients larger than 0.10 s−1 (Fig. 8). At NWG17/33,NWG19, and NWG26/34, these Moho depths are respectively 40,34 and 36 km (Fig. 9).

A previous receiver functions study in this area (Svenningsenet al., 2007) finds a clear Moho at 38–40 km depth close to stationNWG17/33, complying with our estimate. The refraction seismicstudy of Stratford and Thybo (2011b) locates the Moho discontinu-ity at a depth of 36 km close to station NWG19, in agreement withour estimate of 34–36 km. We infer however mantle velocities thatare 0.3–0.4 km/s lower than the velocities in Stratford and Thybo(2011b), and crustal velocities that are 0.1–0.2 km/s lower.

4.2.2. Oslo GrabenWe have three stations located within the Oslo Graben: NWG25

is located at the southernmost border of the graben, NWG32 is rightin the middle, and NAO01 is at the northernmost end of the riftstructure (light green triangles in Fig. 1). From refraction seismics(Stratford and Thybo, 2011a) we know that the extent of the OsloGraben is wider in the lower crust than what appears at the surface.Thus, the permanent station KONO is located outside the surfaceexpression of the Oslo Graben but above a middle and lower crustheavily affected by the rifting (Stratford and Thybo, 2011a).

We did not invert the data at the two stations KONO and NWG32

since their receiver functions show a strong variation with backazimuth, indicating that the assumption of flat isotropic layers isnot satisfied. Probably this is due to a complicated rift structure,and possibly also due to a mantle reflector below the graben, at

194 M.L. Kolstrup, V. Maupin / Precambrian

Fig. 9. Moho depth map. Coloured circles show Moho depth. Grey triangles showthe location of stations with a gradual crust-mantle transition. Black crosses showthe location of stations where the joint inversion was not performed. Black linesdepict main geological units and boundaries from Fig. 1. (For interpretation of thero

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eferences to colour in this figure legend, the reader is referred to the web versionf this article.)

pproximately 65 km depth (Stratford and Thybo, 2011a). The pro-les at the two remaining stations are similar (Figure A.4).

The velocity profile from station NAO01 (Fig. 6 plot E) showsncreasing velocities in the upper 10 km, approximately constantelocities in the mid crust, and gradually increasing velocities inhe lower crust. At a depth of 34 km a velocity jump from approx-mately 4.05–4.30 km/s defines the Moho discontinuity. At NAO01

e can compare our results with two previous receiver functiontudies: Svenningsen et al. (2007) finds a Moho depth of 32 km andttemöller and Midzi (2003) a Moho depth of 34–36 km.

.2.3. Bamble and Telemark blocksStations NWG31 and NWG27 are located in the Bamble Block,

lose to the south coast. Both models from the Bamble block haveradually increasing velocities in the crust, from approximately.4 km/s at the surface to 3.8 km/s at 32 km depth, and Moho at4 km depth (Fig. 6 plot F and Figure A.5).

The receiver functions from NWG31 show some variation withack azimuth, but models inverted using different parts of theata set are very similar, showing a clear Moho discontinuity at4 km depth. This Moho depth also fits well with the 33 km Mohoepth found by Sellevoll and Warrick (1971) in the southern end ofheir Fedje-Grimstad refraction profile. The cause of the variationith back azimuth could be that NWG31 is overlying the contact

etween the Telemark and Bamble Blocks, where the Telemarklock is underthrust below the Bamble block (Lie and Husebye,

994).

Station NWG24 is located in the Telemark block, northwest ofhe Bamble block. The velocity model from NWG24 is very simi-ar to the model from the neighbouring station NWG27, the only

Research 238 (2013) 186– 198

difference being that the Moho discontinuity is slightly deeper atabout 36 km depth (Fig. 6 plot G). Since stations NWG31, NWG27,and NWG24 only have phase velocity periods between 3 and 30 s(Table 1), we do not regard the differences in mantle velocitiesbelow 45–50 km as significant.

NWG20 is located on the Mandal-Ustaoset Shear Zone betweenthe Telemark and the Hardanger-Rogaland Block (see Fig. 1). Thestation has relatively few receiver functions and the 4 invertedstacks contain only 2 or 3 traces each (Table 1). The velocity modelfrom NWG20 has a more gradual transition from crust to mantlevelocities than the other models in the Telemark and Bamble Blocks(Fig. 8 and Figure A.5). The maximum velocity gradient has a valueof 0.08 s−1 and is found at a depth of 34 km. Mantle velocities ofmore than 4.4 km/s are reached at a depth of 40 km. The gradualcrust-mantle transition is dictated by a broad Moho Ps conversionin the lower frequency receiver functions (Gaussian width factor of1), which is split into several small pulses in the higher frequencyreceiver functions (Gaussian width factor of 2.5).

4.2.4. Hardanger-Rogaland blockThe velocity models for this block are shown in Figure A.6 and

the representative profile from station NWG30 is displayed in Fig. 6plot H. Stations NWG22, NWG29 and NWG30 are located at thecoast, while stations NWG28, NWG23/35 and NWG21/ODD1 arelocated progressively inland (red triangles in Fig. 1). The coastalmodels NWG22 and NWG30 have clear Moho discontinuities atdepths of 30 km and 28 km, respectively. The inland stations haveMoho discontinuities at greater depths: 32 km for station NWG28,and 36 km at stations NWG23/35 and NWG21/ODD1. The Mohodepth estimate at ODD1 complies with the estimate of 34–38 kmof Ottemöller and Midzi (2003). It should be noted that stationNWG28 is located close to the Mandal-Ustaoset shear zone, butthe velocity model has a well defined Moho discontinuity veryunlike the gradual crust-mantle transition further north at stationNWG20.

Station NWG29 is located south of Stavanger on a fragment ofCaledonian cover. The velocity model from this station is very dif-ferent from the other velocity profiles in the Hardanger-RogalandBlock. The crust-mantle transition is gradual below 28 km depth,and in the mid crust we have a prominent low velocity zone. Thelow velocity zone is dictated by a trough preceding the Moho Psconversion in all receiver functions from NWG29. The velocity pro-file from NWG30 (Fig. 6 plot H) also shows a tendency of a lowvelocity zone in the mid crust, but with a much smaller amplitude.

4.2.5. Southern Western Gneiss Region and Caledonian NappesThe area between the Hardanger-Rogaland and Telemark blocks

and the Western Gneiss Region is covered by the Caledonian JotunNappes in the west and the Osen-Røa Nappes in the east. In thiscentral area of southern Norway several crustal studies have beenperformed recently (Svenningsen et al., 2007; Stratford and Thybo,2011b). Stations NWG12, NWG13, NWG15, NWG16 and NC204 arelocated in this area (dark green triangles in Fig. 1) and the veloc-ity profile from station NWG13 is shown in Fig. 6 plot I, while allprofiles are shown in Figure A.7. Receiver functions from stationNWG15 show a strong back azimuthal variation, possibly relatedto a Moho offset previously imaged by Svenningsen et al. (2007).Hence, we do not invert for velocities at this station.

All velocity profiles from this area have a clearly defined Mohodiscontinuity and no prominent crustal low velocity zones. Atstation NWG13 velocities increase from 3.95 to 4.42 km/s at depthsof 38–40 km, with a maximum gradient that places the Moho dis-

continuity at a depth of 38 km. This is slightly deeper than theestimate of 36 km in Stratford and Thybo (2011b), but the crustalvelocities are quite similar. Differences between the crustal veloc-ities in our profile and the cross section of Stratford and Thybo

M.L. Kolstrup, V. Maupin / Precambrian Research 238 (2013) 186– 198 195

est of

(vNcAad

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NsvfFav

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Fig. 10. Velocity profiles from stations in the Western Gneiss Region, southw

2011b) are less than 0.1 km/s until depths of 36 km, but mantleelocities depart by 0.2–0.3 km/s. The velocity model from stationWG16 displays a clear Moho discontinuity at 42 km depth, whichonforms with the estimate of 40 km in Svenningsen et al. (2007).t stations NWG12 and NC204 crustal velocities and Moho depthsgree well with Stratford and Thybo (2011b), but mantle velocitiesepart by 0.2–0.3 km/s.

Compared to the velocity profiles from the Telemark andardanger-Rogaland Blocks the inland stations with Caledonianovers have a deeper Moho (Fig. 9) and a more simple layeredtructure with monotonously increasing velocities.

North of the Caledonian covers, stations BER, NWG14, FOO,WG07, NWG09 and MOL are located in the Western Gneiss Region

outh of the Sveconorwegian Front (violet triangles in Fig. 1). Theelocity profile from station NWG14 is displayed in Fig. 6 plot Jor comparison with other areas, while all profiles are shown inig. 10. The profiles from the Western Gneiss Region are more vari-ble than the inland profiles, and several of the profiles have lowelocity zones in the middle and lower crust.

At station BER (Bergen) we find a Moho depth of 30 km, com-lying with a previous estimate from receiver functions of 29 kmSvenningsen et al., 2007) and 30 km from refraction seismicsStratford and Thybo, 2011b). At the permanent station FOO the

oho is found at approximately 34 km depth, conforming withhe estimate of 32–36 km in Ottemöller and Midzi (2003). StationsWG07 and MOL are located 40–60 km from the coast in the mosteavily deformed part of the Western Gneiss Region (Hacker et al.,010), and show Moho depths of 38 and 38–40 km depth, respec-ively. These estimates conform with receiver function studies inhe area (Ottemöller and Midzi, 2003; Svenningsen et al., 2007), butre deeper than the estimate of 34 km at MOL from refractions seis-ics (Stratford and Thybo, 2011b). Compared to the shallow Moho

epths at the southern and western coasts of southern Norway, wend a quite thick crust in these parts of the Western Gneiss Region.lso at station NWG09, a bit further inland, we find a deep but sharpoho at 40 km depth.Between stations NWG03 and MOL (Fig. 1) we have an inter-

sting transition from a gradual crust-mantle transition to a ratherell defined Moho discontinuity (Figs. 8 and 9). At both stationse also have some variation of the receiver functions with back

zimuth. At station NWG03 the Moho Ps conversion is clearer fromhe SW quadrant, and at MOL arrivals at 8 and 9 s change polar-ty with back azimuths from N to E. Despite this, inversions of allub-data sets at MOL were similar with respect to Moho depth and

the Sveconorwegian Front. Locations are denoted by violet triangles in Fig. 1.

character of the crust-mantle transition, and we therefore presentthe velocity profile inverted from all receiver function stacks atthis station. The back azimuthal variation is possibly related toan apparently flat-lying mantle reflector at 46 km depth seen inrefraction seismic studies (Stratford and Thybo, 2011b). Our resultssuggest that the boundary between a crust with a gradual crust-mantle transition and a crust with a sharp Moho discontinuity isclose to these two stations and most likely runs in between them.

5. Discussion

Our results fit overall well with previous geophysical stud-ies in the area (Kinck et al., 1993; Ottemöller and Midzi, 2003;Svenningsen et al., 2007; Stratford and Thybo, 2011b,a; Frassettoand Thybo, 2013) but adds new information, especially on the crust-mantle transition and in the less studied eastern part of SouthernNorway.

In the southwestern part of the study area we infer a sharpMoho discontinuity in the depth range of 28–42 km (see Moho mapin Fig. 9). The general increase in Moho depth from southwest tonortheast, as well as the small crustal root in the middle of south-ern Norway and the thinner crust in the northern part of the OsloGraben, is in good agreement with previous studies, in particularFrassetto and Thybo (2013). We refer therefore to Frassetto andThybo (2013) for a detailed analysis of the isostatic balance of thearea and a discussion on the location and role of the crustal rootwith respect to topography.

In the northeastern part we infer a deep and gradual crust-mantle transition between depths of 40 and 50 km (Figs. 6 and 9).This gradual transition is very consistent from station to stationand had not been clearly seen in earlier studies due to the lack ofproper coverage of seismological stations. It is already suggested inFrassetto and Thybo (2013) but it is more unambiguously seen heredue to the enhanced resolution obtained by combined analysis ofreceiver functions and surface wave dispersions. The crust-mantletransition is therefore qualitatively different in these two areas sep-arated by a boundary striking northwest. The two areas are clearlyseen in the map of maximum crust-mantle velocity gradient inFig. 8.

In the Archean and Proterozoic Fennoscandian Shield major

boundaries between tectonic and accreted domains strike north-west to north (e.g. Gaál and Gorbatschev, 1987; Korja andHeikkinen, 2005). The boundary zone found in our study (Figs.8 and 9) follows this Proterozoic northwest strike and not the

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outhwest strike associated with the Caledonian orogen (e.g.oberts, 2003). This suggests that the boundary zone is a Protero-oic feature and that the deep crustal structure has been preserveduring the Caledonian orogeny.

.1. The gradual crust-mantle transition

The area with gradual crust-mantle transition coincides withhe Eastern Segment (Section 4.1.1), the northern part of the

estern Gneiss Region and of the Caledonian Nappes (Sec-ion 4.1.2). This part of the Western Gneiss Region is not affectedy the Sveconorwegian orogeny (Tucker et al., 1990; Austrheimt al., 2003), but the crust below the Caledonian Nappes mostikely experienced Sveconorwegian deformation. The part of theastern Segment studied here comprises granitoids of the Transs-andinavian Igneous Belt (TIB) (Högdahl et al., 2004) of latealeoproterozoic age (1.71–1.65 Ga Söderlund et al., 1999; Bingent al., 2008; Andersen et al., 2009a; Lamminen et al., 2011hat were metamorphosed during the Sveconorwegian orogenySöderlund et al., 1999). A genetic link has been suggested betweenhe northern half of the Eastern Segment (TIB) and the Hustadgneous Complex in the Western Gneiss Region (Gorbatschev,980; Austrheim et al., 2003). The Hustad Igneous Complex is

ocated just north of the estimated Sveconorwegian Front andhow similar characteristics in age (1.65 Ga, Austrheim et al., 2003)nd geochemistry as the northern Eastern Segment (Persson et al.,995). In both areas we infer a gradual crust-mantle transition withery similar lower crustal S-wave velocity structures.

Högdahl et al. (2004) suggest that the TIB was generated byarge-scale magmatic underplating and reworking of Svecofennianrust in an Andean type subduction zone. The late TIB2 and TIB3ubduction zone (1.71–1.66 Ga) had an approximately northwesttrike with subduction towards the northeast. They infer that theIB magmatism was succeeded by the Gothian arc-related pro-esses forming a more juvenile crust in the Idefjorden TerraneAhäll and Larson, 2000). The southern part of the TIB has been

apped by near-vertical marine reflection profiles in the Balticea, providing detailed images of the crust and uppermost man-le (BABEL Working Group, 1993; Korja and Heikkinen, 2005).elow the granitoid TIB rocks large mushroom-shaped plutons atepths between 20 and 60 km are associated with very weak Mohoeflections. These plutons are interpreted as the heat source of theIB magmatism (Korja and Heikkinen, 2005). The deep and dif-use Moho discontinuity associated with the massive magmaticnderplatings in the Baltic Sea could be similar to the gradual crust-antle transition we find below the TIB in the Eastern Segment.The extent and continuation of the TIB below the Caledonian

appes is not known exactly, but from the associated magneticnomalies the TIB should continue from the Eastern Segment toentral Norway and Lofoten (Högdahl et al., 2004; Olesen et al.,010). The area in the northern Western Gneiss Region wheree find a gradual crust-mantle transition is therefore not consid-

red part of the TIB, though there are similarities in age (1.65 Ga,ustrheim et al., 2003) and geochemistry between this part of theestern Gneiss Region and the TIB rocks in the Eastern Segment

Persson et al., 1995). Due to the strong similarity in lower crustalelocity structure, along with the genetic link in age and geochem-stry, we suggest that the entire area with thick crust and lack of

clear Moho discontinuity has been formed by the same type ofagmatic processes in the late Paleoproterozoic.Large-scale magmatic underplating similar to the late

1.71–1.66 Ga) TIB magmatism (Högdahl et al., 2004), as imaged inhe Baltic Sea (BABEL Working Group, 1993; Korja and Heikkinen,005), can explain the presence of a thick crust. However, it cannotxplain the lack of a clear Moho discontinuity, as even a high

Research 238 (2013) 186– 198

density lower crust (e.g. Al-Kindi et al., 2003; Stratford and Thybo,2011a) will give a larger velocity gradient than the one we observe.

Leech (2001) describes how shortage of fluids can lead to aneclogitization of the lower crust which is only partial, e.g. a small-scale mixture of metastable gabbro with mafic eclogite. Lowercrustal rocks typically have observed Vp velocities of 7.0 ± 0.3 km/s(Rudnick and Fountain, 1995), while mafic eclogites at lower crustaldepths have Vp and Vs velocities of more than 8.0 km/s and 4.5 km/s,respectively (Manghnani et al., 1974). A boundary between a nor-mal lower crust and a fully eclogitized mafic lower crust wouldtherefore be defined as the seismic Moho (Baird et al., 1995),whereas an only partially eclogitized lower crust would make amore diffuse boundary between crust and mantle (Leech, 2001).

Massive magmatic underplates that are only partially eclogi-tized below 40 km depth can explain the deep crust and gradualcrust-mantle transition. A deep crustal root that is only partiallyeclogitized has lower densities than a fully eclogitized root andwill not tend to delaminate, which prevents extensional collapse(Leech, 2001). Arrested orogens with preserved collisional struc-tures and deep, diffuse Moho discontinuities have been describedin the South Ural Mountains (Berzin et al., 1996; Knapp et al., 1996)and also in the Precambrian Trans-Hudson orogen in North America(Baird et al., 1995). The deep and gradual crust-mantle transitioninferred in this study could therefore be an indicator of an arrestedorogenic development.

5.2. The sharp Moho discontinuity

The region where we infer a sharp and shallow Moho disconti-nuity coincide with the area that was intensively deformed duringthe Sveconorwegian orogeny (Bingen et al., 2008). Bingen et al.(2006, 2008) claim that the Sveconorwegian Belt underwent post-collisional extensional collapse in its final stages, after ca. 0.97 Ga.Slagstad et al. (2012) describe a non-collisional tectonic settingfor the Sveconorwegian orogen and suggest back-arc extension insouthern Norway with melting of the lower crust after 0.99 Ga anddelamination or convective thinning of a thickened mantle rootbetween 0.95 and 0.92 Ga. Leech (2001) claims that post-collisionalextensional collapse is driven by delamination of eclogitized lowercrust. This removes the high velocity lower crust, giving a clearMoho discontinuity (Leech, 2001). Another cause of a flat andwell-defined Moho can be flow of the lower crust during exten-sional collapse driven mainly by mantle lithosphere delamination(Nelson, 1992). In all cases it seems likely that the sharp and shal-low Moho discontinuity is related to extension in the late stages ofthe Sveconorwegian orogeny.

Extensional collapse following the Caledonian orogeny mighthave caused a similar resetting of the Moho, but it cannot explainhow the inferred boundary zone strike northwest, approximatelyperpendicular to the Caledonides. Our results therefore suggestthat the present-day S-wave velocity structure of the crust-mantletransition in our study area is related to magmatic and tectonicprocesses in the late Paleoproterozoic. In the Western GneissRegion this velocity structure must have survived the Caledonianorogeny. These findings lend support to the tectonic model ofHacker et al. (2010) in which the Western Gneiss Region is thrustover relatively undeformed Fennoscandian crust during Caledonianexhumation.

6. Conclusions

We present 42 S-wave velocity profiles of the crust and upper-most mantle in southern Norway and western Sweden, estimatedfrom joint inversions of teleseismic P-receiver functions andRayleigh wave phase velocities.

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In the northeastern part of the study area, S-wave velocitiesncrease gradually from lower crustal ones at about 40 km depth to

antle ones at about 50 km depth. A possible explanation for thisradual crust-mantle transition is partial eclogitization of massiveagmatic underplates.In contrast, the S-wave velocity profiles in western and cen-

ral southern Norway show mostly a sharp and well defined Mohoiscontinuity at depths between 29 and 42 km. The sharp andell-defined Moho discontinuity can be a consequence of Neopro-

erozoic extension in the late stages of the Sveconorwegian orogen.The boundary zone between a sharp Moho discontinuity and

gradual crust-mantle transition has a strike approximately per-endicular to the Caledonides, suggesting that it survived thealedonian orogeny.

cknowledgements

This work has been done in the framework of the ESF EURO-ORES TOPO-EUROPE Program 07-TOPO-EUROPE-FP-014. Wehank Andy Frassetto and Hans Thybo for providing us with receiverunctions from the MAGNUS and DANSEIS networks, the NORSARrray and from permanent stations in Norway and Sweden. Like-ise, we thank Andreas Köhler for providing us with surface waveispersion curves. MAGNUS waveforms were recorded with theobile KArlsruhe BroadBand Array (KABBA) of the Karlsruhe Insti-

ute of Technology, Germany as well as with permanent stations ofhe NORSAR array and the Norwegian National Seismological Net-ork. Financial support for the MAGNUS experiment was provided

y the Universities of Aarhus, Copenhagen,Karlsruhe and Oslo asell as NORSAR. We are grateful to Robert B. Herrmann for mak-

ng Computer Programs in Seismology freely available. This workmproved from discussions with Tom Andersen, Arild Andresen,ars Eivind Augland and Roy-Helge Gabrielsen, and from the con-tructive reviews of two anonymous referees. We thank Micheleeremans for map data. Figures have been prepared using theeneric Mapping Tools (Wessel and Smith, 1998).

ppendix A. Supplementary data

Supplementary data associated with this article can be found,n the online version, at http://dx.doi.org/10.1016/j.precamres.013.10.004.

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