Geophysical Journal International

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1 Geophysical Journal International Geophys. J. Int. (2010) 182, doi: /j X x Sedimentary and crustal structure from the Ellesmere Island and Greenland continental shelves onto the Lomonosov Ridge, Arctic Ocean H. Ruth Jackson, 1 Trine Dahl-Jensen 2 and the LORITA working group 1 Geological Survey of Canada, Dartmouth, Nova Scotia, Canada 2 Geological Survey of Denmark and Greenland, Øster Voldgade 10, 1350 Copenhagen K, Denmark. tdj@geus.dk Accepted 2010 March 19. Received 2010 March 19; in original form 2009 April 27 SUMMARY On the northern passive margin of Ellesmere Island and Greenland, two long wide-angle seismic reflection/refraction (WAR) profiles and a short vertical incident reflection profile were acquired. The WAR seismic source was explosives and the receivers were vertical geophones placed on the sea ice. A 440 km long North-South profile that crossed the shelf, a bathymetric trough and onto the Lomonosov Ridge was completed. In addition, a 110 km long profile along the trough was completed. P-wave velocity models were created by forward and inverse modelling. On the shelf modelling indicates a 12 km deep sedimentary basin consisting of three layers with velocities of , and km s 1. Between the km s 1 and km s 1 layers there is a velocity discontinuity that dips seaward, consistent with a regional unconformity. The km s 1 layer is interpreted to be Palaeozoic to Mesozoic age strata, based on local and regional geological constraints. Beneath these layers, velocities of km s 1 are correlated with metasedimentary rocks that outcrop along the coast. These four layers continue from the shelf onto the Lomonosov Ridge. On the Ridge, the bathymetric contours define a plateau 220 km across. The plateau is a basement high, confirmed by short reflection profiles and the velocities of km s 1. Radial magnetic anomalies emanate from the plateau indicating the volcanic nature of this feature. A lower crustal velocity of km s 1, within the range identified on the Lomonosov Ridge near the Pole and typical of rifted continental crust, is interpreted along the entire line. The Moho, based on the WAR data, has significant relief from 17 to 27 km that is confirmed by gravity modelling and consistent with the regional tectonics. In the trough, Moho shallows eastward from a maximum depth of km. No indication of oceanic crust was found in the bathymetric trough. Key words: controlled source seismology; continental margins: divergent; oceanic plateaus and microcontinents; hotspots; large igneous provinces; arctic region. GJI Geodynamics and tectonics 1 INTRODUCTION The Lomonosov Ridge is a remarkable linear bathymetric high extending across the Arctic Ocean (Fig. 1). The Ridge can be followed from north of the Lincoln Sea to near the North Pole and across to the Siberian continental margin. It is a mere km wide and about 1700 km long. The Ridge rises from ocean depths of greater than 4000 m to as shallow as 400 m below sea level. The Ridge is of major scientific interest because it divides the Arctic Ocean into two unequal sized basins, the Eurasia and Amerasia basins, and its structure is significant for understanding the origin of both regions. The Lomonosov Ridge has been described as a double-side continental margin (Jokat et al. 1992). LORITA Working Group: Deping Chian, John Shimeld, Thomas Funck, Isa Asudeh, Dave Snyder. On the Eurasia Basin side of the Lomonosov Ridge, spreading on the Gakkel Ridge created typical lineated seafloor spreading anomalies (Heezen & Ewing 1961; Vogt et al. 1979; Srivastava & Tapscott 1986). The spreading severed the Lomonosov Ridge as a continental fragment from the Eurasia continental margin (Wilson 1963; Weber & Sweeney 1985; Jokat et al. 1992). However, there are variations in the morphology along the strike of the Ridge (Cochran et al. 2006) and there are no wide-angle seismic reflection/refraction (WAR) profiles available from the Pole to the Lincoln Sea to confirm its continental velocity structure. Magnetic depth-to-source calculations on the North American end of the Ridge indicate shallow sources, possibly due to abnormally thickened and elevated oceanic crust that formed in Late Cretaceous to Paleocene (Kovacs & Vogt 1982). Plate reconstructions based on lineated magnetic anomalies in the Eurasia Basin match the parallel shape of the bathymetric contours C 2010 The Authors 11

2 12 H. R. Jackson et al. Figure 1. Overview map. The depth contours (500 m, 1000 m, 1500 m and 3500 m) are from Jakobsson et al. (2000). Marked in various colours are the projected WAR (LW) and drift seismic (LD) data reported in this report; in black previous data. The coloured dots are the individual positions of all shots (large dots) and receivers (solid coloured lines). The acquisition of data took place on four separate days. All data on the inner line (yellow) were acquired in 1 day, whereas data on the outer line were acquired on 2 days separated by 2 days of snowstorm (green and blue). All data on the cross line (pink) were acquired on 1 day. For modelling, the data were projected onto two lines (the NS and X line) marked in grey. of the Lomonosov Ridge and the conjugate Eurasia margin (Fig. 1). Reconstructions (Srivastava & Tapscott 1986; Gaina et al. 2002; Brozena et al. 2003) restore the Lomonosov Ridge adjacent to this margin. Plate reconstructions of the position of the Lomonosov Ridge relative to its termination adjacent to the polar margin vary. Some models leave the Ridge connected to the North American and Greenland margins (Lawver & Scotese 1990) whereas others require a strike-slip fault along the continental margin in the Lincoln Sea up to magnetic anomaly chron 13 (Lawver et al. 2002; Brozena et al. 2003). However, all the reconstructions fix the Ridge to the North American Plate from chron 13 to present. The origin of the Amerasia Basin is enigmatic with many different models suggested for its origin (Tessensohn & Piepjohn 2000) but is generally considered to be created by seafloor spreading processes in Late Jurassic to Cretaceous time (Sweeney et al. 1982). The consequences for the various models for development of the Amerasia Basin side of the Lomonosov Ridge are that it is either a rifted or a strike-slip margin (Cochran et al. 2006); therefore, data on the Ridge s structure and relationship to the continental margin would be valuable to confirm or eliminate various models. The nature of the transition from the Lomonosov Ridge to the polar margin of North America and Greenland is not well known (Jackson & Gunnarsson 1990; Kristoffersen & Mikkelsen 2006) due to the difficulty of collecting seismic data in the ice infested Arctic Ocean. This region is crucial to understanding the development of the Arctic Ocean, both prior to and during seafloor spreading in the Eurasia Basin. The location of the two WAR profiles (Fig. 1), one line along the Ridge onto the continental margin, and the other line perpendicular to it across the bathymetric trough, was planned to give insight into the crustal structure at this important intersection. 2 BATHYMETRY, GEOLOGICAL AND GEOPHYSICAL SETTING The location of the WAR profiles and of the short seismic reflection line is remote and difficult to access due to the ice and weather conditions. Therefore, a limited regional database exists. Here, we describe the available information that is critical for constraining the new data.

3 Lomonosov Ridge to Ellesmere Island/Greenland Bathymetry The North-South WAR profile crosses the continental shelf from the Lincoln Sea onto the Lomonosov Ridge. The bathymetry along the lines is from the International Bathymetric Chart of the Arctic Ocean (IBCAO) grid (Jakobsson et al. 2000). The Lincoln Sea extends about 200 km (Fig. 1) offshore with water depths between 200 and 400 m. The bathymetry of the shelf is gently undulating (Dawes 1990) except near the mouth of Nares Strait where a narrow 500 m deep channel exists. At the edge of the Lincoln Sea, the depth increases to 2250 m in a trough about 80 km wide (Fig. 2). Water depths decrease at the southern end of the Lomonosov Ridge to less than 500 m. Here, the Ridge broadens into a plateau 220 km across at the 1000 m contour. North of the plateau from 87 N the Ridge narrows to about 50 km across and deepens slightly, continuing close by the Pole towards Siberia at 86 N as a single blocky feature. The sides of the Ridge are slightly steeper on the Amerasia flank than on the Eurasia side (Cochran et al. 2006). Near 160 W the Ridge bends sharply, paralleling the shape of the Gakkel Ridge (Brozena et al. 2003) and consists of several rotated blocks (Cochran et al. 2006). 2.2 Geology The polar margins of Canada and Greenland are considered to be passive margins formed by rifting and/or transform motion (Dawes 1990; Forsyth et al. 1990; Sweeney et al. 1990). Information on the continental shelf geology can only be gained from the onshore geology (Fig. 3) and the constraints available from bathymetric and geophysical data. The sole geological samples in the region are from the 15 short sediment cores on the southwestern slope of the Lomonosov Ridge (Kristoffersen & Mikkelsen 2006) that do not provide insight into the type of deeper sedimentary and crustal rocks. Coring and more recently, drilling, has been carried out on the crest of the Lomonosov Ridge, near the North Pole (Blasco et al. 1979; Grantz et al. 2001; Expedition-302-Scientists 2005). International Ocean Drilling Project (IODP) site 302 (Fig. 1 ACEX) penetrated Cenozoic sedimentary strata that were deposited in shallow water (Backman et al. 2008). An older sample obtained from a piston core from the Eurasian flank of the Lomonosov Ridge contains 250 Ma old zircons. The only known source for zircons of this age is the Eurasia shelf. The provenance of the zircons supports the hypothesis that the Ridge is a continental sliver rifted from the Eurasia margin. 2.3 Potential field data Regional (Verhoef et al. 1996; Glebovsky et al. 1998) and local (Forsythet al. 1994) magnetic surveys are available that are useful in extrapolating onshore geology into the offshore. Irregularly shaped positive anomalies are associated with Greenland contrasting with the long-wavelength negative anomalies on Ellesmere Island (Figs 1 and 4). The change occurs along the western edge of Nares Strait where the Wegener Fault has been interpreted (Damaske & Oakey 2006). A change in crustal character is also shown by the prominent gravity lows on Greenland lessening parallel to Nares Strait on Ellesmere Island (Fig. 5). The background magnetic signature on the shelf in the vicinity of the NS WAR line is long wavelength and slightly negative, similar to the character of the magnetics onshore northeastern Ellesmere Island (Sweeney et al. 1990). In addition, a short-wavelength low-amplitude magnetic anomaly is observed from the coast of Ellesmere Island across the NS WAR line that then bends eastward (Fig. 4, labelled A); its location, at the coast line, coincidences with the boundary between the Clemens Markam and the Hazen foldbelts basins (Fig. 3). The onshore/offshore structural correlation is also supported by the Alert Geomagnetic anomaly Figure 2. The grey lines indicate the water depth along the NS line and the X line from the IBCAO map (Jakobsson et al. 2000). The large grey dots are the depth at shot positions from the IBCAO map, and the black dots are the depth obtained from the modelling of the WAR data, indicating slightly shallower depths. In the Lincoln Sea individual depth soundings were taken along part of the inner line shown in small light grey dots (Jackson & Dahl-Jensen 2007) with a good match on the inner part, deviating from the IBCAO map as the soundings approach the bathymetric trough. On the north flank of the trough the soundings are in good agreement with the IBCAO map. The table lists the shots where we were able to model a water depth from the WAR data. Shots are labelled for example OS3 for Outer Line Shot 3 and XS2 for Cross Line shot 2.

4 14 H. R. Jackson et al. Figure 3. Geology map. The main part of the map is modified from Escher & Pulvertaft (1995), with added information from Okulitch (1991) at the western part of the north coast of Ellesmere Island. A dashed line marks the boundary. Offshore, the 500 m, 1000 m and 1500 m depth contours (Jakobsson et al. 2000) have been added to outline the position of the Lomonosov Ridge. The WAR lines discussed in this paper, as well the Shelf 92 and IG WAR lines are marked in black. (Fig. 4) that can be traced across Ellesmere Island and continues offshore as far as 83 N (Niblett & Kurtz 1991). At the seaward edge of the continental shelf at 84 Nbetween 67 W and 80 W (about 75 km southwest of the X line marked with a B on Fig. 4) there is a distinctive 150 km long positive magnetic high with a width of about 50 km. The amplitude and wavelength are similar to that observed on the nearby Alpha Ridge. The bathymetric trough immediately to the north of this anomaly is associated with linear short-wavelength magnetic anomalies that are continuous onto the Lomonosov Ridge (Figs 1 and 4). Based on the magnetic anomalies it is difficult to distinguish the edge of the Alpha Ridge from the Lomonosov Ridge. However, on the gravity map the Lomonosov Ridge is a clear linear feature (Fig. 5). 2.4 Vertical and wide angle seismic profiles Two seismic refraction profiles were collected in a cross pattern on the drifting LOREX ice station in the vicinity of the North Pole (Forsyth & Mair 1984). The velocity-depth section included a 4.7 km s 1 interval interpreted to be sedimentary rocks overlying a 6.6 km s 1 velocity layer, consistent with continental crust. The depth to the Moho was 27 km. This velocity-depth section is similar to that observed on the Eurasian shelf (Forsyth 1978). In addition, the Ridge is reported to be aseismic (Wetmiller & Forsyth 1978) consistent with a continental origin. A single channel seismic reflection profile (Kristoffersen & Mikkelsen 2006) on the west flank of the plateau of the Lomonosov Ridge (Fig. 1, GreenIce) shows a sedimentary section, in some regions thinner than 50 ms two-way traveltime, overlying an undulating acoustic basement. The sedimentary section thickens westward as the water depth increases away from the plateau of the Ridge. Within the sedimentary section there are features interpreted as basaltic flows (Kristoffersen & Mikkelsen 2006). In 2007 a seismic reflection profile was collected about 60 km east of NS WAR profile on the edge of the plateau (Fig. 1 LOMROG07-3). This profile was acquired from the ice breaker Oden supported by the nuclear ice breaker 50 Let Pobedy. A 48 channel digital profile 15 km in length was acquired (Marcussen 2008; Jakobsson 2008). Drifting ice stations collected the only other reflection data between our study area and the North Pole (Ostenso & Wold 1977; Langinen et al. 2009). The seismic reflection profile from Russian ice station NP-28 drifting between the North Pole and the plateau shows a flat-lying sedimentary unit of about 200 m thickness overlying reflections that are variable in intensity and dip (Soper et al. 1982; Langinen et al. 2009). In the same region, ice station Arlis II (Fig. 1) reported a sedimentary section of up to 500 m thickness that is occasionally pierced by acoustic basement (Ostenso & Wold 1977). From ice station LOREX (Fig. 1), the seismic reflection profile shows that the Ridge crest has a saw-tooth appearance covered with thin sediments truncated by scarps suggestive of en echelon fault blocks (Blasco et al. 1979). The reflection profiles were interpreted as diagnostic of continental origin (Ostenso & Wold 1977). In addition, there are several seismic reflection profiles crossing the Lomonosov Ridge between the North Pole and the Siberian continental margin. The sedimentary character is similar to that observed on reflection profiles acquired from the North Pole towards the plateau (Ostenso & Wold 1977; Gramberg et al. 1991). The character of the reflection profiles and velocities from the accompanying sonobuoy data were indicative of a continental origin for the Lomonosov Ridge (Jokat et al. 1992).

5 Lomonosov Ridge to Ellesmere Island/Greenland 15 Figure 4. Magnetic anomaly data (Forsyth et al. 1994; Verhoef et al. 1996; Glebovsky et al. 1998) has been merged into one map (G. Oakey, personal communication, 2008). The WAR lines are marked in heavy straight black lines; the drift seismic line are short black lines outlined in white; IG shows the position of a WAR line near the purple circular magnetic anomaly and associated higher frequency events caused by the HP volcanic rocks; (A) indicates the position of linear magnetic anomaly that stretches from Ellesmere Island toward the Morris Jesup Rise; KW indicates the Kap Washington volcanics; (B) indicates the position of a magnetic anomaly with character similar to others on the Alpha Ridge. The Alert Geomagnetic Anomaly is marked with a dashed line. CFS Alert is marked with a red dot. Depth contours have been added to outline the Lomonosov Ridge the 1000 m contour around the Plateau with a heavier line. For other locations see Fig NEW SEISMIC REFLECTION PROFILE A single channel seismic reflection experiment was carried out from our drifting ice camp during the acquisition of our WAR data (LD, Fig. 1). A 0.1 L sleeve air gun was fired once a minute. Three hydrophones were suspended 7.6 m beneath the top of the ice and a fourth hydrophone was used to measure the source wavelet. The traces were binned with an average of 10 traces per 40 m bin. The air gun ran uninterrupted for 15.6 days and collected data over a circuitous 60.3 km route. The ice camp was initially about 10 km west of the WAR profile. After circling close to this position it moved west. An L-shaped composite seismic section about 27 km long is shown with a number of interpreted horizons (Fig. 6). The profile reveals a thin sedimentary cover on top of a complicated basement structure. The seafloor is horizon 1 underlain by two unconformities (horizons 2 and 3) based on onlapping and truncating geometries within adjacent units. Horizon 4 is a prominent unconformity between units of markedly different reflection geometries. This is the top of acoustic basement; however, coherent dipping reflections are visible in the basement, labelled as unit d. The lack of seismic penetration, even in regions with thin sedimentary cover, indicates that the unit d is either highly cemented (e.g. calcareous), metamorphic or igneous. If calcareous, the dipping units would be due to fault blocks; however, there is no indication of faults. When the seismic profile is viewed in conjunction with the magnetics (Fig. 4), the circular positive magnetic anomaly with radiating linear anomalies suggest that the basement is composed of volcanic rocks. 4 WAR SEISMIC EXPERIMENT 4.1 Acquisition The WAR profiles were collected during April and early May of 2006 using three helicopters and a ski-equipped fixed-wing aircraft to place instruments on and explosives under the sea ice. Collecting WAR data in the ice covered Arctic Ocean requires that the receivers be protected from the cold; takes place in a narrow time window of operations after the sun comes up and before fog limits flying, and is highly dependent on weather and ice conditions. The sea ice in the region of our WAR experiment was many metres thick, with large pressure ridges several metres high and continuously moving. The ice north of Greenland and along Ellesmere Island is the thickest in the Arctic Ocean because the two principle surface currents, the Beaufort Gyre and the Transpolar Drift, transport the ice to this portion of the polar margin. A further complication to the ice motion in our area of operations is that an ice bridge generally forms across Nares Strait that prevents ice moving rapidly along the

6 16 H. R. Jackson et al. Figure 5. Bouguer anomaly map; data from the National Danish Space Centre (Rene Forsberg, personal communication). The 1000 m and 2000 m depth contours are added in white to indicate the position of the Lomonosov Ridge. The WAR lines are in black. strait. The ice becomes highly dynamic when the ice bridge fails. The risk that our explosives and instruments would be swept down the strait was a constant concern. 1 The WAR data were collected on a 440 km long near northsouth line (NS line), and a cross line (X line) of 110 km length (Fig. 1). The lines were acquired in three parts: inner and outer part of the NS line and the X line. The inner part of the NS line was 130 km long, with receivers at a spacing of 1.3 km. There were four shots on the line and two offset shots to the south. The receivers on the 150 km long, outer part of the NS line, overlapped with the inner line by 32.5 km with a receiver spacing of 1.5 km. On the outer line we fired one offset shot to the south, six shots along the line segment with receivers and three offset shots to the north. On the combined NS line we deployed receivers at 183 locations and recovered receivers at 181 locations. On the X line there were 72 receivers spaced about 1.5 km apart over a length of 110 km with 6 shots at 20 km spacing. The shots were either 175 kg or 350 kg, consisting of one or two strings of 10 charges (each 17.5 kg) of Pentolite. Each string was suspended on a 100 m long rope under the ice. The recording instruments consisted of a Marks Products L4, 1-Hz vertical geophone placed directly on the ice and connected to a Nanometrics Taurus digitizer. A detailed account of the acquisition is described in the field report (Jackson & Dahl-Jensen 2007). As the ice moves continuously, each receiver position was logged separately for each shot at shot time. Offsets used in the modelling were calculated on the basis of these positions. For modelling pur- 1 See Appendix S1 (Supporting Information) for acquisition. poses, two great-circle arcs (one for the NS line and one for the X line) were defined onto which all receiver and shot positions were projected (Fig. 1). This was also done for all receivers separately for each shot. The inner line was shot over a short period of time (2 h) with little movement, whereas the outer line was shot in two segments with several days between, resulting in movement of receivers and pre-loaded shots of nearly 16 km during the interruption; hence the larger distances to the projected line on the NS line, mainly in the northern end (Fig. 1; see also Table 1). The X line was shot over a short period (1 h 18 min) of time with little ice motion. 4.2 Method The WAR data were used to obtain 2-D P-wave velocity models for the entire crust and upper mantle along the lines. The P-wave models were obtained using the ray tracing code RAYINVR and accompanying DMPLTSQR inversion and TRAMP amplitude modelling code (Zelt & Ellis 1988; Zelt & Smith 1992; Zelt & Forsyth 1994). For the initial models we used the water depth from the IBCAO grid (Jakobsson et al. 2000). Water depths were then adjusted using the seabed reflection when recorded. For the shots in deeper water, we were thus able to compare the water depths from the IBCAO grid with the depth from the WAR, and found significantly shallower water in the bathymetric trough (Fig. 2). The number of layers in the model was determined from the data on the NS line and carried to the X line. The models were developed using a combination of forward modelling and inversion of velocities and depths of individual layers followed by a comparison of amplitudes on the observed and synthetic data, starting with the uppermost layers. When an

7 Lomonosov Ridge to Ellesmere Island/Greenland 17 Figure 6. (a) Migrated section of seismic data on the line covered by the drifting ice camp during acquisition of the WAR data (LD on Fig. 1). The data have been merged to eliminate the circular motion the camp moved in on the eastern part of the line. (b) Interpreted section with horizons indicated by numbers and seismic units indicated by letters. Horizon 1 is the seafloor, horizon 2, 3 and 4 are unconformities. The units a c are constrained by the unconformities. Unit d is a reflection below the acoustic basement, horizon 4. (c) Track plot showing the geographic orientation of the merged section and the location of the merge point. acceptable solution of the inversion had been found for a layer, the parameters were locked, and the next layer down subjected to inversion. Thus, we developed the models from the top down, using both inversion and forward modelling and amplitude modelling for quality control. Initial quality control of the models was based on a visual comparison of the real and synthetic sections. Where significant differences occurred the model was revised. The final model was determined based on an analytical error analysis of the arrival times and a qualitative comparison of the real and synthetic data. 4.3 Velocity models NS line The P-wave model for the NS line is shown in Fig. 7a whereas in Fig. 7b we have added rays traced through to picked arrivals as well as the reflection points for the reflected phases. The fit between the picks and rays traced through the model for all shots are shown in Fig. 8. The tie with the X line is marked with a white line close to outer shot 4 (OS4) (Fig. 7). The receivers cover the area from OS7 to IS3. On the southern end of the line three layers with low velocities are identified (Fig. 9). The three upper layers have velocities of , and km s 1 (Fig. 7a) and are assumed to be sedimentary. The maximum sedimentary infill in the basin is 12 km. The synthetic model for this sedimentary section (Fig. 9) reproduces the energy distribution of the first arrivals indicating that the gradient distribution is appropriate. Beneath the three layers with lowest velocities are arrivals with a velocity range of km s 1. This layer forms a base for the deep basin at the southern end (Fig. 7). We interpret these velocities as metasedimentary rocks because the velocity of common igneous rocks at 200 MPa are generally considered to be >6.0 km s 1 (Salisbury et al. 2003). North of shot OS8, where no seismometers were deployed, the upper velocity structure could not be resolved. Four layers with a combined thickness of 11 km having velocities similar to the southern end of the line were introduced. We adopted the layer definitions from the southern end but have no internal resolution between the four layers. The layers are only delineated as a traveltime delay on two shots (OS10 and 9) (Figs 7b and 8). A prominent feature in the velocity model occurs just north of the bathymetric trough where basement rises up to the seabed or as near as we can determine. The velocity in the upper section of the of 5.9 km s 1 layer (orange) is controlled by diving waves (e.g. Fig. 10) and increases to 6.5 km s 1 at the base of the layer (Fig. 7a). This

8 18 H. R. Jackson et al. Table 1. Position of the shots in latitude and longitude, distance of the shot along the line from south to north, distance that the shot was offset from the base line, time of shot. NS line Shot Latitude ( N) Longitude ( W) Distance along line (km) Distance to line (km) + to west Charge size (kg) Time of shot (UTC) IS April 9 19:20 IS April 9 20:00 IS April 9 19:00 IS April 9 18:00 IS April 9 19:10 IS April 9 18:10 OS April 21 17:11 OS April 21 16:10 OS April 21 15:40 OS April 21 19:15 OS April 19 00:47 OS April 21 17:30 OS April 21 16:50 OS April 21 16:20 OS April 19 01:50 OS April 19 01:00 X line Shot Latitude ( N) Longitude ( W) Distance along line (km) Distance to line (km) + to north Charge size (kg) Time of shot (UTC) XS May 1 23:28 XS May 1 22:50 XS May 1 22:10 XS May 1 22:20 XS May 1 22:40 XS May 1 23:00 layer cannot be identified south of the bathymetric trough and we have let it pinch out in the model. In the north we lack information but have left it in the model. The crustal velocities outside the basement high are in the range of km s 1. The crust is divided into two layers beneath the basement high although the velocity contrast is slight to non-existent at the top of the lower crust (Fig. 7a). This division is necessary to explain the observed gradient changes between the upper and lower crust. The depth of the diving waves in the crust controls the information we have on the gradients within it (Fig. 11b). Diving rays at about 15 km depth and at a range of 350 km indicate subtle changes in the crustal velocities that are continued northward. On the north end of the line, unreversed diving waves reach a maximum depth of 16.5 km that also constrain the top of the lower crust beneath the high. The Moho depth varies substantially from 20 to 27 km. The base of the crust is defined along the NS line with P m P arrivals from 70 to 420 km range (Fig. 11b). South of 431 km the Moho deepens; we have no WAR data in this area and the deepening Moho is based on gravity data (see further). The high amplitudes on the P m P are an indication of a high velocity contrast at the Moho. The amplitude modelling (Fig. 10) supports continuing the low gradients that we use for the crust based on diving waves that occur only in the upper portion of the layer on the NS line. A test with higher velocities in the lower crust failed. Inversion with DMPLSTSQR returns to a model with lower velocities within a few iterations. Furthermore, the shape of the P m P traveltime curve is affected by the velocity in the lower crust and it is matched by the position of our calculated arrivals (Fig. 10). We find strong apparent post-critical P m P energy on several shots, and very pronounced on OS6, 7 and 8 (Fig. 12). We are unable to model this consistently with the 2-D model and suspect this energy is due to 3-D effects X line The X line model is 110 km long, and the P-wave model is shown in Figs 7c and d. The fit between the picks and rays traced through the model for all shots are shown in Fig. 8. The tie with the NS line is marked in white. The combined thickness of sedimentary layers and the depth to Moho in the velocity model ties well with the NS line, with the crustal layer thickening to the west. The subdivision of the sedimentary section into three layers of km s 1 carried over from NS line, but is poorly defined on the X line due to a lack of diving waves on the cross line alone (Fig. 13). This is due to the deeper water and fewer arrivals at small offsets. In contrast, the km s 1 layer (yellow) is sampled by numerous diving rays so the velocities are well determined (Figs 7c and 11c). The topography on the interface at the top of the 5.5 km s 1 layer is defined by these diving rays (e.g. Fig. 13). The velocities in the upper crust ( km s 1 ) are also defined by diving waves (Fig. 11d). There is no direct evidence for a division into an upper and lower crust on the X line; the geometry is carried over from the NS line and represents a modest increase in velocity from 6.5 to 6.6 km s 1. The depth to Moho decreases from a maximum of 20 km at 40 km range from the western end to 16 km at the eastern end. 4.4 Model resolution and uncertainty As described earlier we have chosen to use the RAYINVR and accompanying DMPLTSQR inversion and TRAMP amplitude modelling code, thus applying a forward modelling and linear inversion to a non-linear problem. This limits our ability to get quantitative estimates of model parameter uncertainties and resolution and a sense of non-uniqueness (Zelt & Smith 1992), as choice of

9 Lomonosov Ridge to Ellesmere Island/Greenland 19 Figure 7. (a) The velocity model on the NS line and (b) with all ray paths used for modelling traced through. The blue dashes are reflection points. The white line marks the tie with the X Line. Panels (c) and (d), as (a) and (b) for the X line.

10 20 H. R. Jackson et al. Figure 7. (Continued.) parameterization and our input of aprioriknowledge (sparse as it is in this study) influences the modelling. Comparison with other modelling approaches and choice of modelling stategy are described in Zelt (1999) and Zelt et al. (2003). Below we present information and estimates on model uncertainties and resolution. The lines were acquired with explosives that produce high signalto-noise arrivals, where phase coherence can confidently be followed across the sections. However, due to the receiver spacing of 1.5 km there are few observations at short offsets, and the consequence of this is less resolution in the shallow velocity layers. In general, the horizontal resolution of the model is determined by the shot and receiver spacing, being half of the greater distance. Our shot spacing is about 30 km on the NS line and 20 km on the X line so the horizontal resolution near the surface, where the diving waves are steep, is 15 and 10 km, respectively. As the rays travel deeper in the crust the paths are more horizontal and the resolution decreases. The errors on the position of the layer boundaries also increase with depth. A sensitivity analysis was done by altering the depth of the interfaces. The layer boundaries above the Moho are known to about ±0.5 and ±1.0 km for the Moho, our results are similar to general uncertainty ranges of km s 1 for velocities and from ±0.2 to 2 km for Moho depths (Zelt & Smith 1992). The ray coverage (Figs 7b and d and 11) gives a good indication of the areas in the models where the velocities and boundaries are well resolved. From this it is obvious that the central part of the NS line has better resolution than the edges. The intersection between the two models occurs in the bathymetric trough, providing further confidence in our models here because we have ray coverage at right angles. Although there is some ray coverage at the northern end of the NS line from OS10 to OS7 and at southern end of the line from IS3 to IS1, the shots are unreversed due to the lack of receivers north of OS7 and south of IS3 (Fig. 7a). The rms and χ 2 values for individual phases are summarized in Table 2. Here the rms error between the picks and the modelled traveltimes as well as the normalized χ 2 values, are listed. The normalized χ 2 is based on pick uncertainties of ms depending on the quality of the individual picks. Pick uncertainties are illustrated graphically in Fig. 8. For some phases the χ 2 is substantially larger than the optimum value of 1. The data are of high quality, and the high χ 2 values indicate that the model is under-parameterized; that is the true nature of the crust is more complex than our model allows for (Zelt & Smith 1992). One possible cause of the high χ 2 values is that the pick uncertainty is too low and that errors in the near-surface propagate with depth. In particular, spot soundings in this section of the Arctic Ocean are scarce so detailed bathymetric maps are not available.

11 Lomonosov Ridge to Ellesmere Island/Greenland 21 Figure 8. For each shot the upper panel shows the rays traced through the velocity model (NS line for shots IS1-6 and OS1-10 and X line for XS1-6) to picked and identified arrivals. The lower panel shows the match between the picks (vertical line; the length is the assigned uncertainty of the pick) and the calculated travel time through the velocity model (black line). The time axis is reduced with a velocity of 6.5 km s 1.

12 22 H. R. Jackson et al. Figure 8. (Continued.) As discussed earlier, the shots and receivers on the NS line deviate up to 10 km from the base line used to model the data (Fig. 1 and Table 1). OS5 for example is significantly off line (Fig. 10) and the fit between observed and calculated traveltimes is poor. This we deem to be a 3-D effect of complex crustal geometry NS line resolution We recorded a total of 1622 traces on the NS line, and used 1571 traveltime picks for the modelling. An additional 167 picks were assigned to assumed offline arrivals. The refraction in the upper crust (Table 2) is an example of a phase with a high χ 2 value. The structure the seismic waves travel in is probably highly non-2-d, and combined with the shots and receivers being scattered over a width of 16 km around the projected model line, we attribute the large χ 2 values to our attempt to model 3-D structures in a 2-D model. The sedimentary and metasedimentary structures (Figs 11a and b the yellow and green layers) in the southern end of the line are resolved both by reflections and diving waves. The depth of the sedimentary basin is constrained by reflected events. The basement high is defined by many diving waves in addition to a few reflections from the top of the layer. The mid-crustal boundary is required to explain diving waves from unreversed shot 0S10 that provides constraints between 50 and 150 km and diving waves from OS1, IS4 and IS1 that constrain the velocities in the lower crust at ranges of and km, respectively. Furthermore, to properly match all of the P m P phases across the model a two

13 Lomonosov Ridge to Ellesmere Island/Greenland 23 Figure 9. Shot IS3. In panel (a), data plotted with reduced time of 6.5 km s 1. The distance scale at the bottom panel is distance along the line. Panel (b), same as panel (a) but with picked phases in red and arrival time of rays traced through the model in green. Ps 1 is the diving wave through the top layer of sedimentary rocks, Ps 2 through the second layer and so on. Psr 3 (on the figure Ps3r) is the reflection of the base of layer three; that is the base of the sedimentary basin. Pg is the diving wave through the crystalline crust, P m P the Moho reflection and P n the diving wave in the mantle. (c) Synthetic data (Zelt & Ellis 1988) to enable comparison of amplitudes. The picks at km have no rays traced through the model, however, the amplitude modelling matched the picks well. (d) Section of the velocity model with ray paths traced rays. Only that part of the line where receivers where active for this shot is shown 1 = water layer; 2 4 = sedimentary layers; 5 = metasedimentary layer; 6 = upper crust; 7 = lower crust; 8 = mantle. For velocity scale see Fig. 7. layer crustal structure is needed (Fig. 7). Moho depth is controlled by reflections for the larger part of the model, clearly outlining the root under the basement high X line resolution We recorded a total of 432 traces on the X line, and used 490 traveltime picks for the modelling. The χ 2 values (Table 2) are generally smaller on the X line than on the NS line because the shots and receivers here are on the same line with no offsets. In addition, this line is more along strike with geology. The division of the sedimentary column into three layers was carried over from the intersection with the NS line. The only observed sedimentary arrivals on the X line are diving rays in the middle of the layer. The velocity of the upper crust, that forms the distinct basement high on the NS line, is constrained by diving waves, whereas the definition of the lower crustal layer is included only on the basis of the tie with the NS line. 5 GRAVITY MODEL From the final velocity model for the NS and X lines, 2-D gravity models were created. The velocity models were divided into rectangular blocks, each bordered by boundary or velocity nodes of the velocity model. Each block was assumed to have a constant density ρ, calculated from the average velocity (v) of the block, using an empirical relationship: ρ = v 0.598v v v 4 (v in km s 1, ρ ingcm 3 ), which approximates the curve used (Ludwig et al. 1970). A density of kg m 3 was assigned to the upper mantle (Fig. 14a). Although the gravity modelling approach does not generate unique model results, it can provide a valuable check on our velocity model,

14 24 H. R. Jackson et al. Figure 10. Shot OS5. Panel (a) Data plotted with reduced time of 6.5 km s 1. The distance scale at the bottom panel is distance along the line. Panel (b) same as panel (a) but with picked phases in red and arrival time of rays traced through the model in green. P s1 is the diving wave through the top layer of sedimentary rocks, P wr is the seabed reflection. P ms is the diving wave in the metasedimentary layer. P g is the diving wave through the crystalline crust, P m P the Moho reflection and P n the diving wave in the mantle. (c) Synthetic data (Zelt & Ellis 1988) to enable comparison of amplitudes. The high amplitude P m P on the synthetic data is a good match to the data. (d) Section of the velocity model with rays traced. Only the part of the line where the receivers were active for this shot is shown. 1 = water layer; 2 4 = sedimentary layers; 5 = meta-sedimentary layer; 6 = upper crust; 7 = lower crust; 8 = mantle. For velocity scale see Fig. 7. especially the Moho topography. The gravity model is extended beyond the ends of the WAR data to avoid 2-D edge effects. The match between calculated and observed gravity values at the centre of the NS line in the vicinity of the bathymetric trough is within a few mgals. The fit deteriorates to the north at about 100 km range where the basement high penetrates almost to the seabed. On the southern end of the line the match between the calculated and observed gravity values over the sedimentary basin on the continental margin matches the slope of the actual gravity data well. Our calculated gravity anomaly contains longer wavelengths than the observed except for the sharp fluctuations just to the north of the basement high due to a measured change in water depth that may be local. Because water depths are not known in detail, this scarcity of soundings may contribute to the misfit. The variation in crustal thickness along the line are consistent with those calculated by Alvey et al. (2008) using gravity inversion techniques for the high Arctic that incorporated a lithospheric thermal gravity correction. From the bathymetric trough to the south for 100 km the calculated and observed gravity values are of similar wavelength and anomalies reach a maximum (Fig. 14a). Here, we observe a thickening in the density layer of 2.25 g cm 3 that is coincident with the seaward edge of the gravity high. A gravity high is a common feature along many Arctic margins; for example to the west along the Canadian polar margin in the Beaufort Sea where the gravity high is associated with a prograding sedimentary wedge of Miocene age or younger (Stephenson et al. 1994). There is also a contribution to the gravity anomaly from the shallowing Moho on our profile. The layer with a density of 2.65 g cm 3 is correlated with metasedimentary rocks of the Franklinian Basin that are observed at the coast of both Ellesmere Island and Greenland (Fig. 3). The shallowing of Moho near 400 km range is required by the P m P

15 Lomonosov Ridge to Ellesmere Island/Greenland 25 Figure 11. (a) NS line: Diving waves in the three layers of sedimentary rocks and the layer of metasedimentary rocks. The reflection points on the interfaces above the crystalline crust are marked in blue the corresponding rays are not marked. (b) Diving waves in the crystalline crust, and reflection points on the Moho. (c) As a for the X line. (d) As b for the X line.

16 26 H. R. Jackson et al. Figure 11. (Continued.) reflections. The deepening of the Moho, at the southern end of the line, is outside the constraints of the WAR data, but is introduced based on the gravity modelling (Fig. 14a) and supported by teleseismic data. Receiver function analysis on data recorded at Alert on Ellesmere Island (Darbyshire 2003) indicates that the crustal thickness at the coast is between 26 and 32 km. On Ellesmere Island crustal thickness varies between 33 and 37 km. Similarly, crustal thicknesses of km were calculated on the west coast of Greenland that increase up to 49 km in the interior of Greenland (Dahl-Jensen et al. 2003). On the North Greenland coast thicknesses of km are determined (Dahl-Jensen 2007). For the X line the calculated and observed gravity curves are similar (Fig. 14b). The densities are slightly different at the tie with the NS line as they average out differently when gradients are converted into blocks of constant densities. 6 DISCUSSION 6.1 Structure sedimentary cover, crust and Moho The velocities and depths determined on the basis of modelling the WAR data (Figs 7a and c) are interpreted in a regional context. Prior to our study it was thought that three major rock units existed beneath the Lincoln Sea: the Cenozoic to Mesozoic Sverdrup Basin sequence, a lower Palaeozoic basin and an attenuated continental basement (Dawes 1990) Sedimentary section The lowest velocities of km s 1 were determined on the NS line and along the Shelf 92 line (Fig. 1). Drilling onshore near the coast of the Arctic Ocean west of Axel Heiberg Island identified a seaward prograding wedge of Tertiary to Upper Cretaceous sedimentary rocks (Meneley et al. 1975). Northwest of Ellesmere Island on the shelf (Fig. 1 near IG) similar velocities were measured and interpreted as Tertiary-Cretaceous strata (Asudeh et al. 1989). Based on the velocities of these layers, their thickening seaward geometry, the associated gravity high and exposed rocks of the Arctic continental terrace wedge (Kerr 1982), we interpret this uppermost unit of our model as belonging to a Tertiary to Cretaceous sequence. Velocities of km s 1 are attributed to the Sverdrup Basin Mesozoic to Upper Palaeozoic sequence (Asudeh et al. 1989; Forsyth et al. 1994; Argyle & Forsyth 1994), or possibly to Mesozoic rocks with similar velocities found offshore in the Mackenzie Delta deposits (Stephenson et al. 1994). To determine if the base of our 9 km thick layer contains Palaeozoic age rocks, the

17 Lomonosov Ridge to Ellesmere Island/Greenland 27 Figure 12. Shot OS8. Panel (a) Data plotted with reduced time of 6.5 km s 1. The distance scale at the bottom panel is distance along the line. Panel (b), same as panel (a) but with picked phases in red and arrival time of rays traced through the model in green. P g is the diving wave through the crystalline crust, P m P the Moho reflection and P n the diving wave in the mantle. The unmatched high amplitude apparent post-critical P m P energy can be explained as an off-line reflection. If we assume a reflection from the suggested volcanic feature on the plateau of the Lomonosov Ridge, and estimate the velocity between shot and the reflector to be similar to the velocity in the basement high found on the NS line, then the modelled reflection is a fair match to the otherwise unexplained energy seen on this shot. Attempts to model it in the 2-D model required low lower crustal velocities (c) Synthetic data (Zelt & Ellis 1988) to enable comparison of amplitudes. The high amplitude P m P on the synthetic data is a good match to the data. (d) Section of the velocity model with rays traced. Only that part of the line where the receivers were active for this shot is shown. 1 = water layer; 2 4 = sedimentary layers; 5 = metasedimentary layer; 6 = upper crust; 7 = lower crust; 8 = mantle. For velocity scale see Fig. 7. geological maps in the region were reviewed. Carboniferous rocks assigned to the Sverdrup Basin outcrop along the coast on northern Ellesmere Island (Okulitch et al. 1990). Plate reconstructions (Fig. 15) place Svalbard near our profile. On Svalbard rocks of Carboniferous age equivalent to the Sverdrup Basin are mapped. The lithological similarities between the Carboniferous succession in the Sverdrup Basin and contemporaneous deposits in northeastern Greenland and Svalbard invite formation-by-formation comparison (Nassichuk & Davies 1980). A basin located between Svalbard and the Sverdrup Basin in the position of our WAR line, prior to seafloor spreading in the Eurasia Basin, would explain the similarities in rock type and age. Therefore, we suggest that it is possible that Palaeozoic strata occur in this unit. This thick section with velocities of km s 1 thins abruptly oceanward in contrast to the overlying younger rocks that thicken seaward. If this thinning is due to faulting, then it affects rocks equivalent to those in the Sverdrup Basin and older and not the younger succession. However, the variations in crustal and sedimentary thickness that we observe are similar to those of the Sverdrup Basin (Sobczak et al. 1990) that are not interpreted to be affected by strike-slip. On the basis of the WAR data alone we cannot distinguish between abrupt changes in crustal layers due to strike-slip faulting or basins created by rifting. Velocities of km s 1 are assigned to Lower Palaeozoic to Upper Proterozoic Franklinian rocks (Argyle & Forsyth 1994). The Franklinian metasedimentary rocks are widely distributed in the

18 28 H. R. Jackson et al. Figure 13. Shot XS1 Panel (a) Data plotted with reduced time of 6.5 km s 1. The distance scale at the bottom panel is distance along the line. Panel (b), same as panel (a) but with picked phases in red and arrival time of rays traced through the model in green. P MS (PMS on the figure) is the diving wave in the metasedimentary layer. P g is the diving wave through the crystalline crust, P m P the Moho reflection and P n the diving wave in the mantle. The structure at the top of the metasedimentary layer (yellow, km) is supported by the pull down in the P m P at the red arrow. (c) Synthetic data (Zelt & Ellis 1988) to enable comparison of amplitudes. The high amplitude P m P on the synthetic data is a good match to the data. (d) Section of the velocity model with rays traced. Only the part of the line where the receivers were active for this shot is shown. 1 = water layer; 2 4 = sedimentary layers; 5 = metasedimentary layer; 6 = upper crust; 7 = lower crust; 8 = mantle. For velocity scale see Fig. 7. Canadian Arctic and underlie the Sverdrup Basin (Trettin 1991). Furthermore, rocks of the Franklinian Basin are exposed on Greenland and Ellesmere Island near the southern end of our profile (Fig. 3). Comparison of our velocity and gravity model across the Lincoln Sea with similar models across the Sverdrup Basin based on geological and seismic reflection control (Sobczak et al. 1990) show remarkable similarities in velocities and geometries of the layers Crust Significantly, the crustal velocity modelled along the entire NS and X lines of km s 1 is similar to those determined on the Lomonosov Ridge near the Pole (Forsyth & Mair 1984). They are also in the range of velocities measured on the margin northwest of Ellesmere Island (Asudeh et al. 1989) and the Beaufort Sea continental margin (Stephenson et al. 1994). Velocities in the range of km s 1 are often associated with rifts (Holbrook et al. 1992) compatible with the location of our measurements on the continental margin. A prominent feature of the velocity model (Fig. 7a) is the basement high that forms the plateau on the southern end of the Lomonosov Ridge and the deep Moho that mirrors it. The upper crust rises to within 1 km of the seabed with a velocity of 5.9 km s 1. Its velocity increases to 6.5 km s 1 at depths of km and its lateral extent broadens significantly. At the surface this body is 23 km across and at mid crustal depths it is about 250 km across (Fig. 7a). Coincident with this feature, linear magnetic anomalies (>200 nt) radiate, bounded by 1000 m bathymetric contours (Fig. 4), suggesting the origin of the 220 km broad plateau is volcanic (Kovacs & Vogt 1982). Furthermore, the two seismic reflection profiles collected on the flanks of the plateau, one by Kristoffersen & Mikkelsen (2006) and the other reported here, indicate dipping reflectors that have been interpreted as volcanic flows.

19 Lomonosov Ridge to Ellesmere Island/Greenland 29 Table 2. Number of shots, no. shots, number of observations, no. of picks, RMS misfit between calculated and picked arrival times, rms, Normalized χ 2 for individual phases. Phase No. shots No. picks rms (s) Normalized χ 2 NS line Seabed refl Sed. refractions Sed. Reflections Refraction metasediments Reflections base metaseds Refraction upper crust/basement high Refraction lower crust P m P P n All phases X line Seabed refl Sed. refractions Sed. Reflections 0 Refraction meta sediments Reflections base metaseds 0 Refraction upper crust/basement high Refraction lower crust 0 P m P P n All phases Figure 14. The velocities from the WAR models were converted to densities (marked in g/cm 3 ) and the gravity response is plotted in grey. Observed gravity data in black are from The Danish National Space Center ( (a) NS line. The deepening of the Moho to 29 km at the southern end the lines is based on the gravity as we have no WAR data here. (b) X line. s = sedimentary rocks ms = metasedimentary rocks Lower crust and comparison with Shelf 92 WAR data Our ray coverage through the lower crust on the X line where the coverage is best indicates velocities from 6.0 to 6.5 km s 1.Onthe NS line the velocity in the lower crust is based on three constraints: the ray coverage in the upper few kilometres of the lower crust, modelling of the shape of curve of the Moho reflection and large amplitude of the Moho reflection. The velocity range based on the above is km s 1. This is consistent with the cross line that intersects it in the region of the trough. We suggest that the velocities are too low to be consistent with underplating. In contrast, near our WAR experiment on the continental shelf and slope a 70 km long WAR profile (Shelf 92) was acquired in 1992

20 30 H. R. Jackson et al. Figure 14. (Continued.) (Argyle et al. 1992; Forsyth et al. 1994) (Fig. 1). They reported velocities of km s 1 for the lower crust. These velocities are typical of margins that are underplated. We have re-modelled the Shelf 92 line (Argyle et al. 1992) using our new velocity models (Fig. 1). 2 The seismic record sections were scanned. We used the velocity model at the western end of the X line as the starting model for the Shelf 92 data, and slightly adjusted the depth and relief of some related boundaries, until the modelled traveltime curves fit the observed arrivals reasonably. The Shelf 92 data clearly displays velocities of km s 1, similar to our shallowest two layers. Beneath these layers a seaward dipping interface separates the shallower velocities from those of km s 1 that were assigned to the upper crust. Velocities of km s 1 are similar to the velocities we determined but labelled (Argyle et al. 1992) as the mid crust. In contrast, we associate and km s 1 layers with sedimentary fill and a metasedimentary layer. Our modelling shows that the 68 km long Shelf 92 line is not of sufficient length to constrain the velocity of the lower crust. There are no diving waves through the lower crust. Wide-angle reflections from the Moho (P m P) are observed, constraining the traveltimes to the Moho but not the lower crustal velocity. Therefore, we conclude that the modelled high velocity ( km s 1 ) lower crust (Argyle et al. 1992) is not well-constrained, and that the much lower velocities of km s 1 found on our longer line are equally valid for modelling the lower crust Moho In our model the velocity in the upper mantle is 7.9 km s 1.The depth to the Moho varies significantly along the NS WAR profile based on our data (Fig. 7b). Here, we consider whether the geolog- 2 See Appendix S2 (Supporting Information) for re-modelling of Shelf 92. ical, geophysical and potential field data support the results of the modelling of the WAR data. From the southern continental edge of the gravity model the Moho climbs towards the coastline at 430 km range (Fig. 14a). The depth to Moho near the coast is 24 km (Fig. 7a), similar to values obtained from teleseismic receiver function analysis at Alert (Darbyshire 2003). P m P arrivals are recorded on the NS WAR from this point northward and suggest that the Moho deepens in km range. Here we have seismic control only on the basement and Moho position (Fig. 7b). However, based on exposures of Sverdrup Basin and Franklinian rocks at the coast it is reasonable to expect these sequences to be present, thus they have been incorporated into the model. Along most continental margins, Moho depth steadily increases shoreward, without a sharp inflection at the shore line. Plate reconstructions for the region postulate a sinistral transform fault here between the North American and Greenland plates from chron 24 to 13 (Srivastava & Tapscott 1986). Large scale transforms such as the San Andreas or Dead Sea fault (DESERT_Group 2004) often, but not always, exhibit Moho topography and abrupt changes in sedimentary and crustal thickness. The shallowing of Moho at this point in the model is consistent with the proposed Nares Strait fault (Fig. 1). From 360 to 300 km range the Moho shallows seawards. The crust and Moho arrivals are well-constrained here. The observed and calculated gravity trends are in close agreement. From about 300 to 200 km range on our model (Fig. 7a) Moho topography varies by only 2 km, gradually deepening seaward. These Moho depths are consistent with the depth on the cross line at 225 km range and similar to those on the Shelf 92 line at a distance of less than 70 km (Figs 1 and 3). From 200 to 160 km Moho deepens and forms a keel beneath the bathymetric plateau with its associated high frequency magnetic anomalies. The gravity model (Fig. 14a) confirms the shape of the Moho here. It is possible that the magnetic anomalies that are observed near the surface may have had a deeper seated origin that has affected the Moho depth. North of the basement high we have limited control (Fig. 11ab) of the sedimentary and crustal structure. There are a few P m P arrivals that indicate the Moho shallows here. The 0 50 km range of our model is where the wide plateau of the Lomonosov Ridge tapers and the water depths increase. A shallowing in the Moho depth at this point is consistent with the changes in bathymetry and narrowing of the Ridge. If the Moho shallows here, as the sparse data indicate, then the depth of the Moho beneath the trough is more representative for the entire Ridge making the plateau anomalously thick, possibly due to volcanic activity. Alternatively, if the thickness of the plateau is average for the entire Lomonosov Ridge, then crustal thinning beneath the trough due to rifting or transform motion must have been focused here during the creation of the Arctic Ocean. The nearest available Moho depth beneath the Lomonosov Ridge is 27 km close to the North Pole (Mair & Forsyth 1982). 6.2 Geological and tectonic implications Origin of the trough The presence of the bathymetric trough might indicate either significant thinning of the continental crust or the presence of oceanic crust. Therefore, we compare our modelled crustal structure with typical oceanic crust. By convention (Christensen & Salisbury 1975; White et al. 1992), oceanic crust is described as having two crustal

21 Lomonosov Ridge to Ellesmere Island/Greenland 31 Figure 15. A plate reconstruction at chron 24 time modified from Tessensohn & Piepjohn (2000). At this time, spreading along the Gakkel Ridge between the North American and European plates is occurring; the Greenland plate moves independently. Inset: Reconstruction at Chron 13 time modified from Soper et al. (1982). From chron 13 time when sea-floor spreading stopped in the Labrador Sea, Greenland moves with the North American Plate and spreading takes place in the North Atlantic. layers plus an overlying sedimentary layer. The shallower crust, layer 2, has a range of velocities with a high velocity gradient and is characterized by extrusive basaltic lavas and dykes formed at a spreading centre. The deeper crust, layer 3, constitutes over twothirds of the crustal thickness and has seismic velocities typical of gabbroic rocks. Velocities at the top of layer 2 are typically 5.04 ± 0.69 km s 1 and at the top of layer 3 are 6.73 ± 0.19 km s 1 (Christensen & Salisbury 1975; White et al. 1992). The velocity of layer 2 is associated with a high gradient and is more variable than that of layer 3.