Author Query Form

Author Query Form Book title: Author: Chapter:

(M35) Arctic Petroleum Geology Y. Kristoffersen 45

The following queries have arisen during copy-editing your manuscript. Please provide an answer in the right-hand column below and on your proofs. Many thanks for your help.

Query No.

Query

Q1

Please confirm the correct authors Rothrock and Maykut (1999) in text citation and Rothrock et al. 1999 in reference list here and elsewhere.

Q2

Please confirm the correct spelling of Tiemann et al. (1982) in text citation and Tieman et al. (1982) in reference list.

Q3

Please confirm the correct spelling of Schlindwein et al. (2007) in text citation and Schwindlein et al. (2007) in reference list.

Q4

Grantz (1993) is not in reference list. Please provide complete publication details to insert in the reference list

Q5

ERS 1998 not in References? Please supply full details.

Q6

Please confirm the correct spelling of author Keynon et al. (2008) in text citation and Kenyon et al. (2008) in reference list.

Q7

Does CTD need to be defined.

Q8

Please confirm the correct spelling of Glebovsky et al. (2004). In text citation and Glebovskiy et al. (2004) in reference list here and elsewhere.

Q9

Please confirm the correct spelling of author Bodganov et al. (1998) in text citation and Bogdanov et al. (1998) in reference list.

Q10

Laxon and McAdoo (1994) is not in reference list. Please provide complete publication detail.

Q11

Please provide one more author name, volume number and page numbers for Dove et al. (2009).

Q12

Engen et al. (2008) is not cited. Please cite in the text.

Q13

Please provide volume number and page numbers in Jakobsson et al. (2008).

Q14

Please provide volume number and page numbers for Langinen et al. (2008).

Q15

Please provide page numbers for Litvinova & Glebovsky (2008).

Q16

Please provide another author’s name and page numbers for Moran et al. (2006).

Response

Q17

Please provide publisher and place of publication for Nansen (1904).

Q18

Please check June 21-11 in Tucker 1989?

Q19

Please confirm the correct spelling of author Keyon et al. (2008) in text citation and Kenyon et al. (2008) in reference list.

1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24 25 26 27 28 29 30 31 32 33 34 35 36 37 38 39 40 41 42 43 44 45 46 47 48 49 50 51 52 53 54 55 56 57 58 59 60 61 62 63 64 65 66 67 68 69

Chapter 45 Geophysical exploration of the Arctic Ocean: the physical environment, survey techniques and brief summary of knowledge YNGVE KRISTOFFERSEN Department of Earth Science, University of Bergen, Alle´gaten 41, N-5007 Bergen, Norway (e-mail: [email protected]) Abstract: The areal extent of the ice covered Arctic Ocean is about four times the size of the Mediterranean Sea. More than 50 years after identification of the main bathymetric features in the deep polar basin, our coverage of modern multichannel seismic data is still limited to what a single ship would complete under open-ocean conditions in about two and a half months. In addition, an estimated one-fifth of the deepwater area is inaccessible to icebreaker surveys with surface-towed equipment. Signal-to-noise levels achieved in multichannel seismic data acquisition from icebreakers are sufficient to identify acoustic basement below more than 3 s (two-way travel time) of sediments in spite of reduced source and receiver arrays. Our current reconnaissance stage in deep basin exploration will at least include the coming decade as surveys by diesel-driven icebreakers are largely restricted by contemporary patterns of open leads in the sea ice cover.

The large water depths encountered during the drift of Fram (1893 –1896) came as a surprise to Nansen, who had brought 200 m of bronze wire, 1200 m of single-cord steel line and 1900 m of hemp line to measure ocean depth. The drift track from the Laptev Sea Shelf to the Yermak Plateau passed over the abyssal plains of the Nansen Basin (Fig. 45.1), and led him to conclude that the North Polar Basin was one large basin (Nansen 1904). Harris (1904), who had studied tides on the north coast of Alaska, observed that the tidal wave arrived from the west and was delayed in comparison to the expected arrival of a wave travelling from the North Atlantic across a uniformly deep polar basin. The delay required a bathymetric obstruction in the form of ‘a tract of land, an archipelago or an area of shallow water’. Fifty years later, Worthington (1953) noticed that the deep water in the Canada Basin was 0.5 8C warmer than Nansen’s observations and postulated a barrier reaching 2000 m water depth. The obstruction turned out to be the submarine Lomonosov Ridge, which divides the Arctic Ocean into two large sub-basins; the Amerasia Basin north of Alaska and Siberia and the Eurasia Basin bordering the European polar margin (Fig. 45.1). The ridge rises about 3000 m above the flanking abyssal plains to a water depth as shallow as 500 m, but generally 1000 –1200 m. Lomonosov Ridge was first discovered in 1948 by Soviet ‘High Latitude Air Expeditions’ followed by more systematic mapping in the subsequent years (Ostrekin 1954). Chukchi Cap was discovered in 1950 by the Soviet ice drift station NP-2 (Hunkins et al. 1962), and Alpha Ridge in 1954 by US ice drift station T-3 (Crary 1954). Publication of the first bathymetric map of the Arctic Ocean by the Soviet Union in 1954 showed a detail which profoundly surprised Western scientist and testified to the magnitude of the postwar Soviet effort (Weber 1983). With the information gathered by the time of the International Geophysical Year 1957/1958, all major submarine features of the polar basin had been identified. Historical and modern datasets comprising soundings from Canadian, Russian and US ice stations, US and British nuclear submarine cruises (1958 –1988), echo soundings from the SCICEX-programme 1993–1999, and data from cruises of icebreakers and research vessels of Canada, Germany, Norway, Russia, Sweden and the United States have now been compiled into a new International Bathymetric Chart of the Arctic Ocean (Jakobsson et al. 2008). The data is available at: http://www. ngdc.noaa.gov/mgg/bathymetry/arctic/ibcaoversion2.html.

Arctic Ocean: the physical environment Sea ice A sea ice cover in the Arctic Ocean is maintained by a strong and stable vertical stratification in the upper 200 m of the ocean, which supresses heat transfer from the warm Atlantic Layer below. Also, the high surface albedo of snow-covered ice (snow c. 0.8 v. ,0.1 for water) limits absorption of heat from incoming radiation and further contributes to a low-energy state. The extent of the sea ice cover varies with the annual cycle by a factor of 2, with a maximum in March and minimum in September, and there is evidence for a long-term decrease of about 6% per decade in areal extent (Johannessen et al. 1999) as well as a decrease in ice thickness (Rothrock & Maykut 1999). Q1 Ice drift. Sea ice drift in the Arctic Ocean is wind-driven with

a first-order pattern expressed by an anticyclonic Beaufort Gyre and the Transpolar Drift forced by a mean high sea-level pressure over the Arctic Ocean. Variations in the forcing on time scales from weeks to a decade are revealed by multivariate analysis of sea-level pressure which transforms the data into uncorrelated orthogonal components. The first principal component which accounts for most of the variability (59%) is called the Arctic Oscillation (Thompson & Wallace 1998). A second principal component (13% of total variance) defined for the winter season has a strong meridional component in the wind and is important for sea ice export (Watanabe & Hasumi 2005). In a negative phase of the Arctic Oscillation (AO), high air pressure at sea-level enhances anticyclonic circulation, the Beaufort Gyre becomes wider (Fig. 45.2a), and an intensified Transpolar Drift removes younger ice faster from the Eastern Siberia and the Laptev Sea margin (Proshutinsky & Johnson 1997; Rigor et al. 2002). During the positive phase, circulation in the Beaufort Gyre is reduced and its centre moved towards Alaska at the expense of a wider Transpolar Drift (Fig. 45.2b). On average, it takes more than 6 years to drift from the Beaufort Sea to the Fram Strait. It takes one year longer during positive AO phases, but ice travels faster from the North Pole to the Fram Strait and induces divergence and production of more thin sea ice over the Eurasia Basin (Rigor et al. 2002).

From: Spencer, A. M., Embry, A. F., Gautier, D. L., Stoupakova, A. V. & Sørensen, K. (eds) Arctic Petroleum Geology. Geological Society, London, Memoirs, 35, 685– 702. 0435-4052/11/$15.00 # The Geological Society of London 2011. DOI: 10.1144/M35.45

686

Dietrich Beaufort et al. (1989) ALASKA Sea Grantz et al. Banks (2009) Isl.

˚

70

Klemperer et al. (2002) Grantz et al. (1998)

75

˚

Basin

Makarov Basin

MJR

Amundsen

Jokat et al. (2008)

East Siberian ev ley Sea e nd e e M idg R Sekretov (2001)

Alpha Ridge

Jokat Lomo noso v et al. (1995)

Gakkel YP

Fram Strait

Fig. 14 a

Fig. 14 b

New Siberian Isl.

Rid

ge

ACEX

Drachev et al. (1999)

Basin Ridge

Laptev Sea

Nansen Basin

˚

SVALBARD

˚

75

60˚

˚

Taimyr Pen.

Kara Sea

Barents Sea

75

Severnaya Zemlja

Geissler & Jokat (2004)

˚

Ellesmere Island

70

GREENLAND

0˚

0˚

SIBIR

Chukchi Borderland

30

0˚

18

˚ 65

Canada

75

70 71 72 73 74 75 76 77 78 79 80 81 82 83 84 85 86 87 88 89 90 91 92 93 94 95 96 97 98 99 100 101 102 103 104 105 106 107 108 109 110 111 112 113 114 115 116 117 118 119 120 121 122 123 124 125 126 127 128 129 130 131 132 133 134 135 136 137 138

Y. KRISTOFFERSEN

12

0˚

500 km Fig. 45.1. Bathymetry of the Arctic Ocean from http://www.ngdc.noaa.gov/mgg/bathymetry/arctic/ibcaoversion2.html with present coverage of published multichannel seismic lines across the surrounding continental margins. Drilling location of Arctic Coring Expedition (ACEX) on Lomonosov Ridge and IODP Sites 910– 12 north of Svalbard indicated by white dots. Locations of seismic sections shown in Figure 45.14 by heavy red lines.

Ice thickness. Model results suggest that undeformed sea ice may

thermodynamically become over 1 m thick during the first year, ,2 m thick after the second year and reach a limit of about 3 m (Maykut & Untersteiner 1971; Vancoppenolle et al. 2006). Sensitivity studies show that the ice thickness is most sensitive to changes in heat flux from the ocean and the insulating effect from incoming radiation if the snow cover is more than 80 cm (Untersteiner 1990). Further increase in thickness is from mechanical thrusting and ridging in convergent zones, most often involving thrusting of younger and thinner ice as refrozen leads close. The sea ice cover is made up of a mosaic of individual floes of dimensions tens of metres to kilometres. Floe perimeters are marked by pressure ridges (,10 m high) or by freeboard steps (,0.5 m) between young and old ice. The sea ice thickness is generally 2– 3 m in the Central Arctic Ocean, 1 –2 m in the Beaufort Sea and Chukchi Borderland area, and around 2 m in the Nansen Basin (Rothrock & Maykut 1999). Any vessel traversing through the pack ice will also meet irregularly dispersed much thicker multiyear flows. Pressure ridges

criss-cross the ice surface, on average 0.5–2 ridges per kilometre (Weeks et al. 1971). Their keels extend three to four times deeper than their surface sail height (Tucker 1989) and their mean total widths cluster around 70 m (Wadhams 1994). The roughness of the ice –water interface has a spectral distribution broadly resembling the roughness of ocean waves (Rothrock & Thorndike 1980). However, for wavelengths ,20 m the spectral density of subice roughness is over 30 dB larger and scattering will affect frequencies above 35 Hz more severely (Dyer 1984).

Ambient noise Arctic Ocean ambient noise levels are strongly related to ice dynamics (Buck 1968; Makris & Dyer 1986). Noise levels build up and decay over periods of at least a day (Kutschale 1969). Drag from wind and currents on ice surfaces of varying top and bottom roughness, frictional forces and inertial motion leading to cracking, shearing and pressure ridging, are all processes which

Colour online= colour hardcopy

CHAPTER 45 ARCTIC OCEAN GEOEXPLORATION

139 140 141 142 143 144 145 146 147 148 149 150 151 152 153 154 155 156 157 158 159 160 161 162 163 164 165 166 167 168 169 170 171 172 173 174 175 176 177 178 179 180 181 182 183 184 185 186 187 188 189 190 191 192 193 194 195 196 197 198 199 200 201 202 203 204 205 206 207

687


Fig. 45.2. (a and b) Principal patterns of ice drift in the Arctic Ocean after Rigor et al. (2002). Isolines indicate the time required for the ice to exit the Arctic Ocean through the Fram Strait during high Arctic Oscillation index (a) and low Arctic Oscillation Index (b). (c) Extensive melt ponds form on the sea ice surface during the melt season. Picture from Makarov Basin mid-August 1996. (d) The seasonal surface melt removes c. 0.5 m of ice and surface objects will pedestal (photograph courtesy of A. Heiberg).

radiate energy into the water over a wide range of frequencies (Fig. 45.3a). Maximum noise energy during low ice activity appears to be centred around 10 Hz with a 5 dB/octave roll-off to 1 kHz (Dyer 1984). The general spectral shape appears unrelated to noise levels varying about 15 dB between quiet and noisy conditions. Also, the noise is quite isotropic at high levels, but partly anisotropic during quiet conditions (Shepard 1979; Chen 1982). Buck (1968) monitored noise level as a function of depth below the ice by placing a reference hydrophone at 91 m depth and raising a second hydrophone in 6 m steps. For any particular frequency, the noise level increased by about 20 dB down to a depth of one-half wavelength and remained essentially constant below this depth (Fig. 45.3b). Buck & Greene (1964) also compared signal-to-noise ratios (frequency .10 Hz) between a hydrophone suspended at 60 m depth and a seismometer embedded in the ice and demonstrate that the ratio for the hydrophone is about twice that of the seismometer (Fig. 45.3c).

Wave propagation The speed of sound in sea ice varies from 3.1 to 3.8 km s21, being highest at low temperatures (Hunkins 1960). The Arctic Ocean water masses differ from lower latitude oceans by relatively small variations in temperature (21.7 to þ1 8C) and relatively large salinity variations (30 – 35 p.s.u.) from top to bottom. For this reason, the associated sound speed profile is primarily controlled by the salinity structure. Water mass properties change most dramatically in the upper 200 m. An uppermost cold and relatively fresh well-mixed surface layer down to 50 m depth is

frictionally coupled to the ice (Untersteiner et al. 1976). The resulting sound speed (Fig. 45.3d, left panel) has a minimum of 1437 m s21 at the top of the mixed layer, a large positive gradient associated with the halocline (50 –200 m depth) followed by a small constant gradient to abyssal depths (Kutschale 1969). Positive velocity gradients result in upward refraction of propagating energy, forming a near-surface wave-guide (Fig. 45.3d, right panel). Arrivals of the direct wave during a sonobuoy measurement will form a curved line with increasing offsets. The turning point for rays at offsets ,30 km will be shallower than about 1200 m depth (Fig. 45.3b). In the open ocean the axis of the wave-guide (or SOFAR-channel) is at about 700 m depth.

Working conditions in the pack ice Return of daylight by mid-March makes it possible to visually inspect the ice surface in a target area from the air in search of ice floes suitable for establishing an ice camp. Air temperatures at this time may be –25 to – 40 8C. Mid-March to the end of May is the optimum time window for aircraft-supported Arctic operations, with often clear skies and a firm frozen surface to land on. The occurrence of low-level Arctic stratus clouds increases between April and June due an increase in cyclonic activity (Serreze et al. 1993). Low clouds imply reduced light contrast of surface features and frequent white-out conditions. The snow surface starts to melt when the mean air temperature passes 21.2 8C and melt ponds (Fig. 45.2c) appear at 858N at about 1 July (Doronin 1970). Freezing starts again as air temperatures fall below zero during the last half of August. Ship operations

688

208 209 210 211 212 213 214 215 216 217 218 219 220 221 222 223 224 225 226 227 228 229 230 231 232 233 234 235 236 237 238 239 240 241 242 243 244 245 246 247 248 249 250 251 252 253 254 255 256 257 258 259 260 261 262 263 264 265 266 267 268 269 270 271 272 273 274 275 276

Y. KRISTOFFERSEN

in the Central Arctic Ocean are generally carried out between mid-July and the end of September. During the annual cycle, a c. 0.5 m thick layer is melted on the surface during the summer period and about the same thickness of new ice added to the underside during the winter. This vertical ‘conveyor’ makes surface structures pedestal (Fig. 45.2d) and transports sediments frozen to the underside of the ice to the surface.

Research platform alternatives There are five principal ways of carrying out research in an ocean covered by perennial sea ice: (1) establish a camp on the ice and drift passively; (2) force your way through with an icebreaker; (3) go under the ice with a nuclear submarine; (4) ride over the ice surface with a hovercraft; or (5) fly over the polar ice pack with an an airplane or satellite. Aero-surveys and satellites are cost-effective for potential field measurements and microwave imagery of sea ice and water, but waterborne sensors are required for acoustic subsurface imaging.

Ice drift stations

Fig. 45.3. (a) Characteristics of the ambient noise level in the Arctic Ocean at 93 m depth modified from Makris and Dyer (1986). (b) Ambient noise levels v. depth in the Arctic Ocean for various frequencies. After Buck (1968). (c) Signal-to-noise ratios of a hydrophone in the water compared with signal-to-noise ratios of a seismometer deployed on sea ice in the Arctic Ocean. After Buck and Greene (1964). (d) Typical sound velocity profile in the Arctic Ocean (left) and ray paths of the water borne energy between a source and receiver. Modified from Kutschale (1969).

Camps on the ice will most often be on multiyear sea ice (2 –4 m thick) and in rare cases on large ice shelf fragments (ice islands) broken off from outlet glaciers on the north coast of Ellesmere Island (Koenig et al. 1952; Jeffries 1992). The ice island T-3 was 13 8 km (Hunkins 1967) and was determined by electrical resistivity measurements to be about 60 m thick (Plouff et al. 1961). Deployments of small temporary camps when the daylight returns by the middle of March are usually by Twin Otter aircraft. Larger camps are supplied by landing larger planes and air drop or low-altitude parachute extraction from C-130 Hercules (Weber 1979; Johnson 1983). A 2 month camp with 20 persons involves moving 40– 50 tons of cargo plus 70– 150 tons of fuel. One or more helicopters are often available in the camp for regional surveys. The start of the melting season is associated with increasing occurrences of low clouds and white-out conditions and back-haul flight operations rarely go beyond the end of May in the Eurasia Basin and a couple of weeks later in the central Arctic Ocean. Ice camps move with the general large-scale sea ice drift pattern modulated by the local wind field. Extended periods (days to months) with unusual passages of smaller pressure systems may, however, result in superimposed excursions of more than 100 km or alternatively periods of little progress. US ice drift station T-3 made four 150 km long traverses across Alpha Ridge under circumstances like this during the summer and autumn of 1968 and 1969 (Hunkins & Tiemann 1977). An ice camp may have primitive facilities, but represents excellent working conditions for multidisiplinary science programmes: ample space, low ambient noise and an opportunity for all participating scientists to fully focus on their respective objectives. Another advantage relative to open ocean conditions is the possibility to maintain large instrument array geometries over time. In general, the ice surface moves as one unit on scales from kilometres to many tens of kilometres, but may suffer episodic local convergence or divergence events forced by shifts in speed and direction of the surface wind (Popelar & Kouba 1983). The consequences of relative ice movements may be significant for cableconnected array elements (Baggeroer & Dyer 1982), but less so for arrays of single self-contained instruments or sensors connected via radio link. Geoscience programmes in ice camps may include navigation, bathymetry, 3.5 kHz chirp echo sounder, sediment coring, seismic reflection and refraction measurements and gravity observations. Magnetic field intensity is best obtained by airborne surveys or a gradiometer configuration due to the time variations


277 278 279 280 281 282 283 284 285 286 287 288 289 290 291 292 293 294 295 296 297 298 299 300 301 302 303 304 305 306 307 308 309 310 311 312 313 314 315 316 317 318 319 320 321 322 323 324 325 326 327 328 329 330 331 332 333 334 335 336 337 338 339 340 341 342 343 344 345

in the geomagnetic field (Coles & Taylor 1990). A regular land gravimeter placed on the ice is sensitive to long-period vertical motion and needs extra damping for convenient readout (Crary & Goldstein 1957). Echo sounder transducers are suspended below the ice through a hydro hole inside a scientific work space. Making a hole through c. 2 m thick ice with a small auger and chainsaw may take a man-day; the building is later put up over the site. The air gun may have a separate hole and is suspended shallow enough (3– 5 m) to ensure air released from each shot is vented through the hole. Accumulation of air pockets below a camp may generate local stress nodes which could attract crack propagation in the event of regional ice dynamics. A small diver compressor (140 l min21) can supply air sufficient for a 0.6 l air gun fired at 50 m shot spacing or about every 4 min during fast drift. Air guns operated at a low repeat cycle may not seal properly in the cold water ( –1.7 8C), but do in general work well. Traces of alcohol in the air line prevent freezing and increased solenoid voltage improves reliability. Alternatively, one may use a sparker source as in the case of the 9 kJ system on T-3 (Hall 1973, 1979). The seismic signals are received by a single hydrophone deployed 30–50 m away from the source and recorded on a laptop. A complete seismic single channel field system based on modern technology may weigh less than 200 kg (Kristoffersen & Mikkelsen 2006). Figure 45.4 shows examples of data recorded from ice station field camps by analogue and digital single channel seismic systems using air gun or sparker source. In spite of the limited

Ice drift station

Single icebreaker

0.3 l Bolt airgun single hydrophone

4 x 4 l G-guns 300 m 24 channel streamer sea ice 1-2 m thick

1.5

2.0

2.0

2.5 s

Lomonosv Ridge north of Ellesmere Island

Ice station T-3 9 kJ sparker single hydrophone

2.5 s

Mendeleyev Ridge

Ice station NP-28 detonators seismometer array

4.5

6.0

5.0

7.0

7.5 6.0 s

s Canada Basin

Amundsen Basin

Fig. 45.4. Samples of single channel seismic data from the Arctic Ocean acquired from ice stations and multichannel seismic data from an icebreaker.

689

source strength, it is possible to define acoustic basement below the main submarine ridges in the Arctic Ocean. The 9 kJ sparker yields a penetration of 2 s travel time below the abyssal plain. Water depths away from the shelf areas are sufficiently large, so interference from water bottom multiples is not important. Russian investigators have operated seismic reflection programmes on eight of their over 30 ice stations beginning with North Pole-13 in 1965 (Gramberg et al. 1991; Sorokin et al. 1999). They have consistently used detonators (three to five) as seismic source fired at the centre of a cross-legged (550 550 or 1150 1150 m) receiving array. Each leg comprised 12 channels recorded on analogue tape and shot spacing was mostly c. 500 m. The analogue traces were later digitized and summed to form a centre trace. The data has good signal-to-noise ratio and an example from North Pole-28 is shown in Figure 45.4 (lower right). We may further exploit the sea ice surface by placing continuously recording sonobuoys offset at right angles to the drift direction and obtaining simultaneous parallel lines (Kristoffersen & Mikkelsen 2006). However, the maximum offset must be less than the critical distance for emergence of a bottom refracted arrival and is thus water depth-dependent. This concept may be expanded to emulate a full 3D survey by aligning detectors and several seismic sources perpendicular to the prevailing drift direction. We may also emulate a multichannel seismic cable by aligning the row of sensors with the average drift direction as carried out from US ice station Fram-4 in 1982 (Kristoffersen & Husebye 1985). Hydrophones from 20 modified continuously recording sonobuoys were deployed through the sea ice at 100 m intervals trailing the main camp (Fig. 45.5a). Hydrophone depth was 60 m and the source was a 2 l bolt air gun fired at 13 m depth. During 34 days of continuous operation, the ice surface moved at of 0– 25 cm s21 (maximum 15 km per day) and firing rates were continuously updated from one to three satellite fixes per hour to maintain a 50 m shot interval (Fig. 45.5c). The ice floe orientation was monitored daily by theodolite sun shots and stayed within +38 of its starting value (Tiemann et al. 1982), but periods of easterly Q2 wind caused significant cross-array drift (Fig. 45.5b). We were forced to operate the air gun at maximum 100 bar instead of 140 bar to reduce leaks and recorded bottom reflections with a signal-to-noise ratio of 33 dB during periods of slow drift (3 cm s21) and 6 dB or less during fast drift (16 cm s21). Vibrations from vortex shedding along the hydrophone cable became significant above a drift speed of 20 cm s21. This experimental setup successfully acquired about 200 km of multichannel seismic data. Normal move-out corrections are small for the case of a short streamer (2 km) relative to the water depth (3.8 km) and we assumed interval velocities representative of a deep ocean environment to generate stacking velocities. Varying degree of array feathering was clearly a problem as common mid-points within a 50 m bin would be scattered over a distance of more than 500 m perpendicular to the track in the worst cases. The consequence is minor over a level abyssal plain, but a strike line over bottom and subsurface layers sloping 58 would imply a travel time difference of 56 ms between zero and 308 array feathering for the far offset trace. Stacking efficiency would, in this case, be degraded unless traces were selected so that midpoints lay on a line perpendicular to profile direction (Levin 1983). Results of the experiment are shown in Figure 45.5d. The advantage of the ice cover to maintain receiver array geometry has also been successfully used for recording microearthquake activity on the Gakkel spreading centre by a sonobuoy array (Kristoffersen et al. 1982) and seismometer arrays (Schlindwein et al. 2007). Q3 Measurements of heat flow through the seabed in the deep Arctic Ocean are carried with thermistors spaced out on a lance or sediment corer as used in the open ocean. A heat flow programme initiated in 1963 by US Geological Survey on US ice station T-3 produced observations at 356 localities within the Amerasia Basin (Lachenbruch & Marshall 1966, 1969). Other measurements

690

(a) Sonobuoy array 20 channels

sea ice thickness 2 m 250 m

100 m

120 cuinch air gun

Array feathering

array drift direction

100°

50

50°

0

0°

−50

-310°

−100

-260°

(c) Ice drift speed

20

20

16 12

wind speed

10

8 4 0

(d)

ice drift speed

cm s–1

10

April

15

20

25

1

May

5

0 9 10

5.0

5.0

6.0

6.0

7.0 l 00

l 12

29 April 1982

Wind speed in knots

wind direction 100

Wind direction

hydrophones

(b)

Two-way travel time (s)

346 347 348 349 350 351 352 353 354 355 356 357 358 359 360 361 362 363 364 365 366 367 368 369 370 371 372 373 374 375 376 377 378 379 380 381 382 383 384 385 386 387 388 389 390 391 392 393 394 395 396 397 398 399 400 401 402 403 404 405 406 407 408 409 410 411 412 413 414

Y. KRISTOFFERSEN

l 7.0 08

l 24

30 April

from ice stations have been made by Canadian (Judge 1980; Taylor et al. 1986) and Russian groups (Lubimova et al. 1973, 1976). Heat flow measurements have been accompanied by sediment coring to obtain estimates of the conductivity.

Icebreakers The Russian nuclear icebreaker Arktika was the first surface vessel to reach the North Pole – a ‘scientific –practical experimental voyage’ (Anonymous, Polar Record 1978). The Murmansk to Murmansk trip was made in August 1977 with an average speed of 11.5 knots. Today, over 40 years later, there are unfortunately still no reports of a nuclear icebreaker ever being used as a primary research platform for geoexploration. The first penetration of a diesel-driven icebreaking research vessel in 1987 heralded a new era in marine geoscientific research in the Arctic Ocean. Polarstern reached beyond 868N north of Svalbard, carrying out a full complement of geophysical measurements and geological sampling except for the use of towed equipment (Thiede 1988). Towing of seismic equipment in heavy pack ice was pioneered in 1988 by Arthur Grantz of the US Geological Survey from US Coast Guard icebreaker Polar Star in the Northwind Ridge area (Grantz et al. 1992). The air gun was suspended below a 1.27 ton steel ball and a two-channel 150 m long hydrophone cable


Fig. 45.5. (a) Cartoon of the array of recording sonobuoys emulating a seismic streamer trailing the Fram-IV camp site along the prevailing drift direction. (b) The array kept its orientation, but drifted sideways in crosswinds. (c) The ice drift followed closely changes in the surface wind. (d) Sample processed section of data acquired over 32 h of drift in the Nansen Basin at the end of April 1982. Modified from Kristoffersen & Husebye (1985).

was fed through a protective hose and effectively towed from the depth level of the air gun (20 m). This arrangement was later expanded to include six air guns (15 l) attached to a triangular frame below the steel ball (Fig. 45.6). The tow line was now an armored umbilical which carried all feeder lines. A single cable for source and hydrophone cable avoided tangling. About 2295 km of seismic reflection data and 28 sonobuoy measurements were made during the 1988, 1992 and 1993 seasons without loss of equipment (Grantz et al. 1993). Polarstern uses a seismic source of eight air guns suspended within a frame for operations in light ice conditions (Fig. 45.6). The first seismic lines across Lomonosov Ridge (Jokat et al. 1992) were acquired in 1991 using one or two air guns towed below a weight, as done by Grantz (1993). Q4 Broken ice of cubic metrre sizes sweeps along the ship’s hull and enters the turbulent wake behind the stern. It is imperative that cables enter the water as close to the stern as possible. However, from time to time the protective hose gets caught behind an ice block and all the gear is forced to the surface. The heavier the gear, the greater the strain on tow cables. Loss of four air guns from Oden in 1996 prompted the use of a depressor (Fig. 45.6) for downward pull and reduced equipment weight (Kristoffersen 1997a). The depressor is easy to handle, but not as effective as a heavy steel ball in keeping the gear close to the stern, which implies that the gear will more often be forced to the surface. Another concern is towing depth. Towing at 20 m


415 416 417 418 419 420 421 422 423 424 425 426 427 428 429 430 431 432 433 434 435 436 437 438 439 440 441 442 Colour 443 online= 444 colour 445 hardcopy 446 447 448 449 450 451 452 453 454 455 456 457 458 459 460 461 462 463 464 465 466 467 468 469 470 471 472 473 474 475 476 477 478 479 480 481 482 483

Umbical

Trumpet Steel ball 1.3 ton

180 m lead-in 150 m long hydrophone cable

Hydrophone cable

Air guns Protective hose

Hydrophone cable

Air guns

Air guns

Fig. 45.6. (Upper) Towing arrangments for seismic sources used during data acquisition in sea ice by Polar Star (Grantz et al. 1993), (middle) by Oden (Engen et al. 2009) and (bottom) by Polarstern (Jokat et al. 1992).

depth means the gear is below the main propeller wash, closer to the stern and less likely to be forced to the surface – a ‘survival mode of operation’. The downside is lower seismic resolution from a notch at 40 Hz in the source and receiver spectra created by the delayed surface ghost reflection. Towing with a depressor and at shallower depth gives better resolution, but represents a greater risk for loss of equipment – ‘suicide mode’. The greatest risk is related to the hydrophone cable, particularly in situations where sharp course changes are made. The turbulent wake behind the icebreaker is progressively filled up by blocks of broken ice moving randomly with great force. Once the seismic source is forced to the surface, the entire streamer will follow with a high risk of getting the cable pinched between moving blocks of ice (Fig. 45.7c, d). The inevitable result is torn internal signal wires and loss of hydrophone groups, or, if the ice is under some pressure, loss of major parts of the cable. A practical length of a hydrophone cable behind an icebreaker in heavy ice of less than 500– 300 m is most often used. The cable length compared with water depths greater than 1 km implies that reliable stacking velocities cannot be derived by conventional analysis. Velocity information has to be obtained from from sonobuoys. Useful signals from sonobuoys can be received for offset distances out to about 30 km (Jokat 2003; Engen et al. 2009). Sample shot records (Fig. 45.7a) and the resulting stack (Fig. 45.7b) acquired in heavy ice in Makarov Basin with only 11 live channels out of 24, still yields useful data with about 2 s penetration. Use of a second icebreaker in front of an icebreaking seismic acquisition vessel may in general translate into steadier progress,

691

but only if there is no pressure on the ice. The assistance of Russian nuclear icebreakers has been used in two-ship operations in areas influenced by the thicker ice of the Beaufort Gyre (Jokat 2003; Marcussen 2008). A nuclear icebreaker may keep a straight course in þ3 m thick ice, while a diesel-driven icebreaker operating alone prefers to follow leads. Breaking and ramming intact ice may double fuel oil consumption and is restricted to transitions from one lead to the next aligned with the general direction of travel. Leads tend to have frequent choke points (every few hundred metres) where forward momentum is lost and ramming is required. In this situation, the ship may ease astern (þ50 m) with the seismic gear out and accelerate forward again – otherwise the gear has to be retrieved temporarily and redeployed when the obstruction is passed. Seismic data acquisition in heavy sea ice is characterized by high noise levels, the main contributions being the fundamental frequency and higher harmonics of the ship’s propeller and engines (Buravtsev 1995). The noise levels (rms mbar) during single ship operation in 2– 2.5 m thick ice are three to four times the levels where a clear channel has been made by a lead vessel or in open water (Jokat et al. 1994). Observed noise levels diminished dramatically with offset out to about 280 m from the stern and levelled out for more distant channels. A modern icebreaking research vessel has facilities for a complete suite of geophysical measurements and heavy equipment for geological sampling. Introduction of multibeam swath survey capabilities and high-resolution sub-bottom profiling during the first penetration of Polarstern to 868N north of Svalbard in 1987 (Thiede 1988) represented a huge step forward. From 2007 and onwards, three Western icebreakers (Healy, Oden and Polarstern) routinely collect fundamental seabed information in this manner. The bottom track may be lost during active ramming in heavy ice due to air bubbles in the turbulent water. Important data can, however, be obtained by making the ship turn 1808 in one spot to insonify an area of diameter equal to the swath width (Jakobsson & Marcussen et al. 2008). Underway gravity data is also routinely acquired by Polarstern and Healy.

Nuclear submarines The first attempt to reach the North Pole in 1931 in a diesel-driven submarine was aborted, but the expedition successfully demonstrated the potential of a submarine as a science platform. HU Sverdrup carried out water sampling down to 3000 m depth as well as sediment coring through an open hole in a pressurized compartment of the submarine Nautilus at the ice edge north of Svalbard (Sverdrup 1931). US nuclear submarines Nautilus and Skate recorded the first continuous echosounder profiles across the Arctic Ocean in 1958 and 1959 and outlined the main features of the basin morphology (Dietz & Shumway 1961). During 1997, 25 years (1957 –1982) of soundings collected by US submarines were released to the public (Anonymous 1997). For many scientific disciplines, a submarine research platform would represent the ultimate solution for efficient scientific data acquisition in an ice-covered ocean. The Science Ice Excercise (SCICEX) programme (Fig. 45.8) allowed US scientists access to US Navy nuclear submarines either as ‘accommodation missions’, which allocated time to obtain unclassified data on otherwise classified missions (1993 –2001), or to carry out dedicated scientific surveys in the Arctic Ocean (Edwards & Coakley 2003). A major achievement was implementation of swath mapping capabilities and sub-bottom penetrating chirp sonar attached to the belly of USSN Hawkbill (Chayes et al. 1998). The submarine platform moved at 16 knots below the ice and collected swath bathymetry (width three times the water depth), high-resolution images of the upper þ50 m of sub-bottom sediments, gravity measurements as well as sampling a number of

692

(a)

3

4 (b)

5

5 Two-way travel time (s)

Two-way travel time (s)

484 485 486 487 488 489 490 491 492 493 494 495 496 497 498 499 500 501 502 503 504 505 506 507 508 509 510 511 512 513 514 515 516 517 518 519 520 521 522 523 524 525 526 527 528 529 530 531 532 533 534 535 536 537 538 539 540 541 542 543 544 545 546 547 548 549 550 551 552

Y. KRISTOFFERSEN

6

7

6

7


Makarov Basin (c)

(d)

Healy 2005

water mass properties. The dedicated cruises in 1998 and 1999 (Fig. 45.8) acquired geophysical data along 39.250 km of survey track, primarily over the Alaska continental margin, Chukchi Borderland, Lomonosov Ridge, Gakkel Ridge and Yermak Plateau (Edwards & Coakley 2003). Continued access to nuclear submarines for science has become difficult following a substantial reduction in the US submarine fleet.

Aerosurveys Airplanes such as the P-3 Orion aircraft cruising at 450 km h21 with an endurance of up to 6000 km, are very efficient platforms for regional potential field surveys (Childers et al. 2001). Minimizing the effect of temporal field variations is also a major concern in magnetic field intensity measurements. Rapid field changes are important in high latitudes, particularly in the auroral oval where large areas may have changes of .160 nT h21 (Coles & Taylor 1990). The earliest Soviet reconnaissance aeromagnetic surveys were initiated in 1946 (Demenitskaya & Hunkins 1970). A more systematic effort was undertaken from 1961 onwards and results from the Eurasia Basin were published by Rassokho et al. (1966) and Karasik (1968). Karasik (1974) presented a magnetic anomaly map including the Amerasia Basin. The US efforts began with field campaigns over the entire Arctic Ocean in 1950–1952 (King et al. 1964) and later in 1961–1963 (Ostenso & Wold 1971). These surveys documented a first order pattern with a ‘Central magnetic Zone’ of high amplitudes

Oden 2001

Fig. 45.7. (a) Shot files (low cut at 10 Hz) acquired during Healy cruise 0503 from Makarov Basin illustrating signal levels and (b) the result of processing. (c) Successful seismic data acquisition in progress from Healy in 2005 and (d) a case of difficult ice conditions with the entire seismic cable on top of the ice.

over the Alpha and Mendeleyev ridges and more subdued amplitudes in the Eurasia Basin. Canadian surveys in the 1960s covered over the outer Canadian Arctic islands and the continental shelf (Coles & Taylor 1990). Renewed effort by US Navy agencies in the years 1972 –1978 effectively translated into new understanding of the geological history of the Arctic Ocean; identification of a magnetic lineation pattern in the Eurasia Basin supported Karasik’s (1968) earlier work and tied basin evolution to the Cenozoic opening of the North Atlantic (Vogt et al. 1979), and a set of magnetic lineations in the Canada Basin associated with Mesozoic seafloor spreading (Vogt et al. 1982). An international effort to compile and integrate all aeromagnetic surveys carried out prior to 1990 was spearheaded by Geological Survey of Canada and published in 1996 (Verhoef et al. 1996). An update (Fig. 45.9) has been presented by Kovacs et al. (2002). Missions which included gravity, magnetic field intensity and sea surface heights were flown by the US Naval Research Laboratory during 1992–1999 (Brozena et al. 2003). Track spacing was 11–18 km and the resolution in the gravity signal along-track (22 –34 km) was effectively constrained by the low-pass filter used to extract the signal from the noise (Childers et al. 2001). Survey cross-over errors were ,2 mGal. More limited aero gravity surveys have been flown north of Greenland and Svalbard by Denmark, Germany and Norway. The inventory of Russian gravity data from the Arctic Ocean appears mainly in the form of on-ice gravity measurements made during aircraft landings (Maschenkov et al. 1999).


553 554 555 556 557 558 559 560 561 562 563 564 565 566 567 568 569 570 571 572 573 574 575 576 577 578 579 580 581 582 583 584 585 586 587 588 589 590 591 592 593 594 595 Ubiquitous coverage of the gravity field south of 81.58N is pro596 vided by the satellite gravity field derived from the European 597 Space Agency’s Earth Remote Sensing (ERS) satellite altimeter 598 data. Advanced processing of the full waveform has made it poss599 ible to extract the gravity field over ice-covered areas (Laxon & 600 McAdoo 1998). Although track spacing at 758N is c. 2 km, the 601 size of the radar’s foot print makes it impossible to resolve wave602 lengths ,15 km in the gravity signal (Childers et al. 2001). An 603 example of a comparison between the airborne gravity measure604 ments in the Beaufort Sea region with the ERS-derived field 605 Q5 (ERS 1998) yielded an rms difference of 2.55 mGal. Compilation 606 of all available gravity data is the focus of the Arctic Gravity 607 Q6 Project (Keynon et al. 2008), a special study group under the Inter608 national Association of Geodesy with contributions from the 609 circum-Arctic countries, Finland, Sweden, Iceland, the UK, 610 France and Germany (http://earth-info.nga.mil/GandG/wgs84/ 611 agp/abstract.doc). The data (Fig. 45.10) is available from the 612 website: http://earth-info.nga.mil/GandG/wgs84/agp/agp_08/ 613 ArcGP_GRV.scn. 614 615 616 Hovercraft 617 618 Although a hovercraft can move with relative ease across open 619 water and sea ice of any thickness, its potential as an Arctic logistic 620 alternative remains unexplored (Fig. 45.11). A new hovercraft 621 dedicated to polar research was tested for the first time north of

693


Fig. 45.8. Track lines for science missions by US Navy nuclear submarines. Figure courtesy of B. Coakley.

Svalbard during the summer of 2008 (Hall & Kristoffersen 2009). The craft can carry a payload of 2.2 tons, has sleeping quarters for four persons and is equipped with echo sounder/chirp sonar for high-resolution sub-bottom surveys, compressor and air gun for seismic reflection measurements (0.3 l air gun), EM-probe for ice thickness measurements, CTD and winch (to Q7 500 m) for oceanographic measurements and a sediment corer and winch (1100 m). The craft can, under favourable conditions, travel about 500 nautical miles between fuelling points. Some 3300 nautical miles of track were covered during the first summer (ultimo June –primo September). Its performance in the pack ice was up to expectations and the platform may be used in an ice drift station mode or act as a survey vessel in open leads. Use of hovercraft in conjunction with an icebreaker is expected to greatly enhance the efficiency of the total operation.

Autonomous underwater vehicles Detailed mapping and sampling at a submarine volcanic eruption site on the Gakkel spreading centre have been successfully carried out by two autonomous underwater vehicles (AUV) and a tethered platform, providing real-time high-definition video imagery (Sohn et al. 2008). The AUVs completed nine dives in water depths up to 4062 m. In some cases the vehicles surfaced below unbroken ice, but were located and successfully recovered after the floe had been cut by the icebreaker.

694

Y. KRISTOFFERSEN

622 623 624 625 626 627 628 629 630 631 632 633 634 635 636 637 638 639 640 641 642 643 644 645 646 647 648 Colour 649 online= 650 colour hardcopy 651 652 653 654 655 656 Fig. 45.9. Map of anomalies in total magnetic field intensity over the Arctic 657 Ocean of Kovacs et al. (2002). Track spacing is ,20 km (Childers et al. 2001). 658 659 660 Autonomous data buoys 661 662 Much of the thick sea ice in the Beaufort Gyre north of Canada and 663 Greenland appears inaccessible for surface vessels using towed 664 665 666 667 668 669 670 671 672 673 674 675 676 677 678 679 Colour 680 online= 681 colour 682 hardcopy 683 684 685 686 687 688 Fig. 45.10. Gravity map over the Arctic Ocean compiled by the Arctic Gravity 689 Q19 Project (Keyon et al. 2008). Data from http://earthinfo.nga.mil/GandG/ wgs84/agp/agp_08/ArcGP_GRV.scn. 690


Fig. 45.11. Hovercraft as research platform may conduct surveys in open leads or operate as a drift station depending on ice conditions. (http://www. polarhovercraft.no/).

equipment, and support of manned ice drift stations is costly. Autonomous data buoys recording position, meteorological and oceanographic parameters have been in operation in the Arctic Ocean since 1979, but the ARGOS system can only handle small datasets (Untersteiner & Thorndike 1982). More data-intensive measurements can be transferred via the Iridium satellite system (Double et al. 2006). A prototype autonomous buoy for acquisition of seismic reflection data (Fig. 45.12) has recently been successfully tested north of Svalbard (Meyer et al. 2008). A 4.8 kJ sparker source is fired under GPS control and the data are recorded by a single hydrophone and stored on a flash disk. The first second of sub-bottom reflection is extracted from the dataset and subsequently transmitted to shore as short burst telegrams via the Iridium system. System power is from solar panels and a windmill, which limits operation to the light season of the year – April to August. A prototype data buoy for bathymetry is also under development (Hall 2008).

The Arctic Ocean: state of geophysical exploration A growing recognition of the need to compile and join datasets has, within the last decade, produced a new International Bathymetric Chart of the Arctic Ocean based on historical and modern datasets previously unavailable (Jakobsson et al. 2008), continuing improvements in merging old and new aeromagnetic datasets (Litvinova & Glebovsky 2008) and successful progress of the Arctic Gravity Project in compiling a 50 50 Free Air and Bouguer database (http://earth-info.nga.mil/GandG/wgs84/ agp/agp_08/ArcGP_GRV.scn). Bathymetry and the potential field data are our first-order tools for making meaningful interpretations of the geological evolution of the Arctic Ocean as well as planning future surveys. Next is our need to image sub-bottom layering and geological basement structures. About 16 000 km of multichannel seismic reflection data have been acquired to date in the Arctic Ocean, half of that by icebreaker cruises in 2005, 2007 and 2008. The inventory of seismic reflection data from ice stations is 15 648 km (Kristoffersen & Mikkelsen 2004). It is possible to successfully acquire good-quality seismic reflection data from an icebreaker in .2 m thick ice provided the pack ice field is not under pressure. Experience shows that the thick ice north of NW Greenland and the Canadian Arctic Archipelago is not accessible to surveys with towed equipment (Jokat 2003; Marcussen 2008). Other approaches must be used or advantage taken of situations where huge leads exist for days in the thickest ice as evident from satellite imagery. We will highlight some key points in our state of knowledge of the circum-Arctic continental margins, their marginal plateaus and the submarine ridges in the polar basin.


691 692 693 694 695 696 697 698 699 700 701 702 703 704 705 706 707 708 709 710 711 712 713 714 715 716 717 718 719 720 721 722 723 724 725 726 727 728 729 730 Colour 731 online= 732 colour 733 hardcopy 734 735 736 737 738 739 Fig. 45.12. Autonomous seismic data buoy with a sparker source and recording 740 by a single hydrophone. Data are transmitted in real time as short burst messages 741 via the Iridium satellite system. 742 743 744 Eurasia Basin 745 746 Amundsen and Nansen basins. The active Gakkel spreading centre 747 divides the Eurasia Basin into two sub-basins (Fig. 45.1). Only 748 the western part of the Eurasia Basin is explored by multichannel 749 seismic surveys (Fig. 45.13). Rough basement topography is 750 covered by .2 km of sediments below the respective abyssal 751 plains (Jokat et al. 1995; Weigelt & Jokat 2001; Jokat & 752 Micksch 2004). The Nansen Basin, proximal to the glaciated 753 Barents – Kara Sea margin, is about 300 m shallower and sediQ8 ments are thicker (Glebovsky et al. 2004; Engen et al. 2009). 754 755 Amundsen and Nansen basins are associated with linear, albeit 756 low-amplitude magnetic lineation patterns that are symmetric 757 with respect to the Gakkel spreading centre and correlated with 758 the Cenozoic geomagnetic polarity time scale (Vogt et al. 1979; 759 Glebovsky et al. 2006).

695

Gakkel Ridge spreading centre. The western 1000 km of the

1800 km long Gakkel spreading centre is covered by modern multibeam swath mapping surveys and with the effort of the joint Arctic Mid-Ocean Ridge Expedition (AMORE) in 2001 became one of the best surveyed ridge segments in the world’s ocean (Michael et al. 2003). The effort started with Polarstern in 1987 (Thiede 1988), continued in 1991 (Jokat et al. 1995) and was pursued by an extensive survey using a nuclear submarine (Fig. 45.8) during SCICEX (Cochran et al. 2003). East of 708E, the rift valley is filled with sediments. Melt production appears not to scale down directly with eastward-decreasing spreading rate as expected, but is either high or low within segments along the spreading centre (Michael et al. 2003; Jokat et al. 2003). Svalbard–Severnaya Zemlja continental margin. The margins bound-

ing the length of the Eurasia Basin are considered to be rifted, passive continental margins because the Gakkel spreading centre, located in the middle, is flanked by deep basins associated with ridge-parallel magnetic lineations (Karasik 1968; Brozena et al. 2003; Glebovsky et al. 2004). Six seismic reflection lines cross the continental margin (Fig. 45.1) west of 308E (Baturin et al. 1994; Geissler & Jokat 2004). Basement in this area remains shallow to about 100 km south of the shelf edge. The eastern 1500 km of the margin is covered by Russian potential field data, but unexplored by multichannel seismic reflection measurements. Basement tectonic lineaments along the entire margin trend about normal to the present continental margin (Bodganov et al. 1998; Johansson et al. 2005). Q9 Laptev Sea continental margin. The Laptev Sea continental shelf

(Fig. 45.1) is the continental extension of an active oceanic spreading centre where the present average divergence is 0.57 cm/year (DeMets et al. 1990) and estimated Cenozoic crustal extension is about 300 km (Gaina et al. 2002; Brozena et al. 2003). More than 25 000 km of seismic reflection data have been collected since 1984, suggesting that crustal extension on the shelf has been accommodated by a wide rift bounded to the east by a major detachment fault (Lazarev Fault) extending to the middle crust (Franke et al. 2001). Basement topography below the continental slope is interpreted as a shear zone which represents the transition between localized extension at the Gakkel spreading centre and distributed deformation on the shelf (Drachev et al. 1999; Sekretov 2002). Sediment thickness below the present shelf edge is about 7 km and reaches 13 km within the rift centre of the Ust’ Lena Rift Basin on the shelf (Franke et al. 2001; Sekretov 2002). No deep stratigraphic boreholes exist. Lomonosov Ridge northslope. Four seismic reflection transects and

two seismic refraction profiles extend from Lomonosov Ridge into Amundsen Basin (Fig. 45.13). The slope appears to be formed by a staircase of narrow blocks, and the difference in level between acoustic basement in Amundsen Basin and basement on top of Lomonosov Ridge is about 5 km (Jokat et al. 1995; Butsenko & Poselov 2004; Langinen et al. 2008). Gradients are gentler and the Free Air gravity low and magnetic low associated with the foot of slope are also less distinct between the North Pole and the Laptev margin than between Greenland and the North Pole (Brozena et al. 2003).

Marginal plateaus Morris Jesup. Morris Jesup plateau forms much of the North

Greenland continental margin and is presently the least known structure in the Eurasia Basin (Fig. 45.1). Available data include a traverse by ice station Arlis-2 in 1965 (Ostenso & Wold 1977), a brief visit over its northern tip by Polarstern (Jokat et al. 1995) and aerosurveys of magnetic and gravity measurements (Brozena et al. 2003). The plateau has a steep eastern scarp and more

696

760 761 762 763 764 765 766 767 768 769 770 771 772 773 774 775 776 777 778 779 780 781 782 783 784 785 786 787 788 789 790 791 792 793 794 795 796 797 798 799 800 801 802 803 804 805 806 807 808 809 810 811 812 813 814 815 816 817 818 819 820 821 822 823 824 825 826 827 828

Y. KRISTOFFERSEN


Fig. 45.13. Tracks of seismic reflection data in the Arctic Ocean. White tracks represents ice stations, red lines indicate multichannel data collected by Polarstern, Oden, Healy and Russian vessels, black lines are multichannel surveys by Polar Star. Blue lines represent tracks of Healy and Louis St. Laurant in 2008 (from http:// walrus.wr.usgs.gov/infobank/1108ar/html/1-1-08-ar.nav). Scientific drill sites shown by large white dots.

subdued sloping topography towards west and north as well as a strong east – west trending gravity gradient associated with the foot of slope to the north. The origin of the plateau has been tied to formation of the northeastern part of the Yermak Plateau (Feden et al. 1979; Jackson et al. 1984), or alternatively it may be part of a larger Arctic Cretaceous Large Igneous province (Engen et al. 2009). Yermak Plateau. Yermak Plateau forms two different structural trends, a western NNW trend and a northeastern part striking NE (Fig. 45.1). Aeromagnetic data show two domains – a northeastern part associated with high-amplitude (700 –900 nT) magnetic anomalies in contrast to a rather featureless magnetic field over the western and southern part (Feden et al. 1979). Earlier interpretations of magnetic, seismic refraction and gravity data suggested a purely volcanic northeastern part and a continental crust below the rest of the plateau (Feden et al. 1979; Jackson et al. 1984), including a western boundary formed by a transform related ridge of

massive basalt intrusions (Crane et al. 1982). A grid of multichannel seismic data acquired across the plateau west of 108W (Fig. 45.1) shows variations in basement topography up to 3000 m buried below a more than 50 km wide level crest of the plateau (Jokat et al. 2008). The horst forming Sverdrup Bank is considered an extension of tectonic lineaments associated with the Hornsund Fault Zone along the west coast of Spitsbergen farther south. Also, the northeastern Yermak Plateau may be stretched and intruded continental crust (Jokat et al. 2008; Engen et al. 2009).

Amerasia Basin Canadian Arctic continental margin. Crustal transects across the

Canadian polar margin north of Ellef Ringnes Island are based on seismic refraction measurements and gravity modelling (Sobczak et al. 1986). The northern rim of the Sverdrup Basin was uplifted during Campanian–Maastrichtian with erosional


829 830 831 832 833 834 835 836 837 838 839 840 841 842 843 844 845 846 847 848 849 850 851 852 853 854 855 856 857 858 859 860 861 862 863 864 865 866 867 868 869 870 871 872 873 874 875 876 877 878 879 880 881 882 883 884 885 886 887 888 889 890 891 892 893 894 895 896 897

truncation of Upper Triassic to uppermost Creataceous rocks over the distance from Ellef Ringnes Island to Banks Island (Balkwill et al. 1983). The Arctic Terrance wedge deposited north of the rim and below the present outer shelf comprises Maastrichtian and younger clastic sediments (Sweeney et al. 1990). A seismic reflection and refraction transect in the Beaufort Sea –Mackenzie area and gravity modelling suggest a rifted passive margin with a narrow transition zone in the eastern part and a broad area of transitional crust towards the west (Cook et al. 1987; Dietrich et al. 1989; Stephenson et al. 1994; Lane 2002). Arctic margin of Alaska. A single deep seismic reflection line

crosses the northeastern Chukchi Shelf (Fig. 45.1) at 1658W (Klemperer et al. 2002) and acoustic basement is not observed anywhere else in the oceanic domain along the continental slope of Arctic Alaska (Grantz et al. 1994). Regional uplift starting in Early Cretaceous (Valanginian) accentuated as well as eroded a structural high, the Barrow Arc, which extends from Yukon to NE (1638W) Chukchi Shelf (Mickey et al. 2002). The north flank of Barrow Arch represents a hinge line with a dramatic increase in sediment thickness (.10 km) above a Lower Cretaceous Unconformity as well as crustal thinning (Saltus & Bird 2003). The crest of Barrow Arch subsided and a thin condensed Hauterivian sediment section accumulated above an unconformity considered to represent the breakup unconformity at a passive rifted continental margin (Mickey & Haga 1987; Grantz et al. 1990). The Alaskan margin has no associated linear magnetic field signature which could reveal contrasting continental and oceanic crustal domains (Kovacs et al. 2002). Also, in spite of exceptionally thick sediment accumulations along the margin, the free air gravity anomaly is of lower amplitude and more

697

discontinuous than observed along both rifted and sheared North Atlantic continental margin segments. Chukchi Borderland. The Chukchi Borderland (Fig. 45.1) appears

as an assemblage of north-trending linear fault blocks. Gravity modelling suggests that the crustal root of the borderland extends down to 27 km depth (Hall 1990). The geomagnetic field appears smooth over Northwind Ridge and Northwind Plain, but high and uniform over Chukchi Cap (Kovacs et al. 2002). Two depth-transects of piston cores up the steep eastern slope of Northwind Ridge recovered rocks which represented every Phanerozoic system except Devonian and Silurian and document the probably continental nature of the ridge (Grantz et al. 1998). East Siberian Sea continental margin. The continental margin

includes a shelf area up to 800 km wide (Fig. 45.1). Multichannel seismic transects exist north of De Long Island (1608E) (Franke et al. 2001; Sekretov 2001), and a summary of a refraction section has been presented by Grikurov et al. (2003). Thin sediments cover basement over De Long High, increasing to 8 km below the shelf break and over 9 km below the continental slope in 2000 m water depth. The crust is about 7 km thick below the lower continental slope. The geological structure associated with the junction of Mendeleyev Ridge with the continental margin is not known. The East Siberian continental margin appears to have the characteristics of a rifted margin. Lomonosov Ridge –Makarov Basin margin. The Amerasia Basin

facing the upper slope of the Lomonosov Ridge shows a prograding wedge (Fig. 45.14) below the early Cenozoic planar

Fig. 45.14. (a) Seismic section across Mendeleyev Ridge acquired by Healy in 2005 and (b) over Lomonosov Ridge by Polarstern in 1991.

698

898 899 900 901 902 903 904 905 906 907 908 909 910 911 912 913 914 915 916 917 918 919 920 921 922 923 924 925 926 927 928 929 930 931 932 933 934 935 936 937 938 939 940 941 942 943 944 945 946 947 948 949 950 951 952 953 954 955 956 957 958 959 960 961 962 963 964 965 966

Y. KRISTOFFERSEN

unconformity in the central part of the ridge (Jokat et al. 1992; Kristoffersen et al. 2004). Distinct linear bathymetric expressions which include the Marvin Spur and buried basement segments extending towards the Siberian margin are present at the foot of slope accompanied by co-located gravity and magnetic anomaly trends (Jokat 2005; Cochran et al. 2006). Chaotic reflection segments are observed in four crossing multichannel profiles. This tectonic trend has been associated with a major shear zone postulated in models of the Amerasia Basin formed by clockwise rotation of an Alaska –Chukota block (Green et al. 1986; Grantz et al. 1998; Cochran et al. 2006). Canada Basin. Canada Basin is the largest of the polar sub-basins

(Fig. 45.1) and the estimated thickness of predominantly terrigenous sediments varies from .6 km in the western part to more than 12 km towards the Beaufort shelf (Grantz et al. 1990). The relative contribution of hemipelagic sediment input can be assessed from the thickness of the sediment drape (0 –7 –0.8 km) covering the high-standing Alpha and Mendeleyev ridges to the north of the basin and unaccessible to turbidite deposition (Dove et al. 2009). Plate boundaries during evolution of the Canada Basin remain elusive (Grantz et al. 1998). The gravity field shows little variation except for a low-amplitude, but distinct linear trend running approximately north –south from north of the Mackenzie Delta to the foot of Alpha Ridge (Laxon & Q10 McAdoo 1994). Weak magnetic amplitude anomalies are colinear with this feature, which is suggested to represent an extinct spreading centre (Taylor et al. 1981; Kovacs et al. 2002). The intriuging possibility of the deep Canada Basin being floored by oceanic crust associated with this fan-shaped zone of weak linear magnetic anomalies and by extended continental crust elsewhere has been forwarded by Grantz et al. (2011). The sea ice is thickest north of Arctic Canada and the northeastern part of Canada Basin is not accessible for multichannel seismic surveys from icebreakers (Fig. 45.13). Submarine ridges – Lomonosov Ridge. Lomonosov Ridge reaches

water depths of 500–1000 m and divides the polar ocean into two sub-basins (Fig. 45.1). The ridge is 1750 km long and 40– 70 km wide in the central part and 150 –200 km at both ends. The structure appears to be formed by an ensemble of en echelon fault blocks slightly oblique to the main trend (Jokat et al. 1992; Cochran et al. 2006). The internal block stratification or structure remains unresolved. The postulated origin as a sliver rifted off the Mesozoic Svalbard – Severnaya Zemlja margin (Wilson 1963) is since supported by the presence of a prograding wedge on its south side (Jokat et al. 1992). Foresets are truncated by a planar unconformity and scientific drilling has documented ridge elevation at or near sea-level in the early Cenozoic (Moran et al. 2006). On the Eurasia Basin side, the rifted continental margin is formed by a steep staircase pattern of narrow blocks. Only the central part of Lomonosov Ridge (858300 –888N, 1508E) and the end north of Canada/Greenland (848300 – 888N) show peneplained surfaces below a younger sediment drape, indicating subaerial exposure, erosion and subsidence. Submarine ridges – Alpha and Mendeleyev ridges. Alpha and Mende-

leyev ridges form a 250–800 km-wide rise reaching water depths shallower than 2000 (Fig. 45.1). About 4000 km of new multichannel seismic data have been acquired over its accessible part (west of 1508W) by icebreaker cruises in 2005 and 2008 (Fig. 45.13). The ridges are covered by a 700– 800 m thick drape of hemipelagic sediments and acoustic basement is characterized by laterally varying, internal bright reflection segments and subbasement velocities in the 3.5 –5 km s21 range (Dove et al. 2009). Long-range refraction experiments on both ridges show velocity –depth relation distinctly different from a global average velocity –depth functions representing a wide range of continental settings (Forsyth et al. 1986; Christensen & Mooney 1995;

Lebedeva-Ivanova et al. 2006). The nature of acoustic basement as well as the seismic velocity of its deeper structure suggest a volcanic origin of the ridges.

Scientific drilling A proposal for scientific drilling in the central Arctic Ocean (Proposal 533) matured from events which included the seismic documentation acquired in 1991 of a c. 450 m thick hemipelagic drape on top of Lomonosov Ridge (Jokat et al. 1992), a stationkeeping test in drifting sea ice by Oden in 1991 (Kristoffersen et al. 1992), an unsuccessful attempt to sample sediments on the ridge by shallow drilling (Kristoffersen 1997b) and transition to an Integrated Ocean Drilling Program which embraced the concept of mission-specific platforms. An ice-strengthened drilling vessel assisted by two icebreakers penetrated 420 m of Cenozoic sediments unconformably overlying Campanian (c. 80 Ma) shallow marine sands (Moran et al. 2006). The section has a stratigraphic gap of 26 Ma (44 –18 Ma). Scientific highlights include documentation for bi-polar emergence of ice prior to 40 Ma and enhanced bi-polar ice growth at c. 14.5 Ma. The Eocene section contain 4 –14% organic carbon.

Concluding remarks Camps on drifting sea ice and aerosurveys were the principal platforms during the first 50 years of Arctic Ocean exploration. The sea ice drift pattern is irregular and camp programmes were not optimized for sub-bottom exploration. New aerosurveys can improve definition of the gravity and magnetic fields, but the firstorder features are known. The pioneering acquisition of multichannel seismic data in heavy sea ice in 1988 (Grantz et al. 1992) marked the beginning of a new era in Arctic Ocean exploration. The other milestone was diesel-driven icebreakers reaching the North Pole in 1991 (Fuetterer 1992). These new opportunities have unfortunately not been fully realized because of half-hearted efforts to tailor equipment for handling and ice protection of seismic gear in the turbulent wake of an icebreaker, a consequence which directly translates into considerable waste of expensive icebreaker time. This issue needs attention. The efficiency of an icebreaker may be further enhanced by an accompanying hovercraft serving as an independent seismic vessel or mobile ice drift station depending on ice conditions. Our present inventory of seismic reflection data falls short of providing substantial clues on complex issues like the tectonic evolution of the Amerasia Basin. Nevertheless, the reflectivity and reflection geometries of hemipelagic deposits on elevated submarine ridges provide important constraints on the palaeoceanography of the basin and define targets for future scientific drilling (Moran et al. 2006; Dove et al. 2009). The author has enjoyed the benefit of joint seismic field operations in the Arctic Ocean with W. Jokat, B. Coakle and J. Hopper and numerous discussions with A. Grantz. The spirit of the captains and crews of Polarstern, Oden and Healy contributed to successful multichannel seismic operations under sometimes challenging conditions. GMT software (Wessel & Smith 1991) was used to generate some of the figures.

References Anonymous. 1978. Icebreaker voyage to the North Pole, 1977. Polar Record, 19, 67 –68. Anonymous. 1997. 25 Years of Arctic sub bathymetric data declassified. Witness the Arctic, Chronicles of the Arctic System Science Research Program, 5, 16. Baggeroer, A. B. & Dyer, I. 1982. Fram 2 in the Eastern Arctic. EOS, Transactions, American Geophysical Union, 63, 217 –221.


967 968 969 970 971 972 973 974 975 976 977 978 979 980 981 982 983 984 985 986 987 988 989 990 991 992 993 994 995 996 997 998 999 1000 1001 1002 1003 1004 1005 1006 1007 1008 1009 1010 1011 1012 1013 1014 1015 1016 1017 1018 1019 1020 1021 1022 1023 1024 1025 1026 1027 1028 1029 1030 1031 1032 1033 1034 1035

Balkwill, H., Cook, D. G. et al. 1983. Arctic North America and northern Greenland. In: Moullade, M. & Nairn, A. E. M. (eds) The Phanerozoic Geology of the World, II. The Mesozoic. Elsevier, Amsterdam, 1– 31. Baturin, D., Fedukhina, T., Savostin, L. & Yunov, A. 1994. A geophysical survey of the Spitsbergen margin and surrounding areas, Mar. Geophysical Research, 16, 463 – 484. Bogdanov, N. A., Khain, V. Yu. et al. 1998. Explanatory Notes for the Tectonic Map of Kara and Laptev Seas and Northern Siberia (Scale: 1:2.5 mill.) Institute of the Lithosphere of Marginal Seas. Russian Academy of Sciences, Moscow, Special Publications. Brozena, J. M., Childers, V. A. et al. 2003. New aerogeophysical study of the Eurasia Basin and Lomonosov Ridge: implications for basin development. Geology, 31, 825 –828. Buck, B. M. 1968. Arctic acoustic transmission loss and ambient noise. In: Sater, J. E. (ed.) Arctic Drifting Stations. Arctic Institute of North America, Calgary, 427 – 439. Buck, B. M. & Greene, C. R. 1964. Arctic deep-water propagation measurements. Journal of the Acoustical Society of America, 36, 1526– 1533. Buravtsev, V. Yu. 1995. Features of streamer noise during marine seismic explorations under ice conditions. Oceanology, 35, 435 – 437. Butsenko, V. V. & Poselov, V. 2004. Regional features of the seismic configuration of the sedimentary cover in the deep Arctic basin and possibilities of its paleotectonic interpretation. In: Avetisov, G., Basov, V. A., Piskarev, A. L., Progebitsky, Yu. & Trukhalev, A. I. (eds) Geological– Geophysical Characteristics of the Lithosphere in the Arctic Region. Proceedings of Ministry of Natural Resources of the Russian Federation, Russian Research Institute for Geology and Mineral Resources of the Ocean. VNIIOkeangeologia, St Petersburg, 5, 141– 168. (Translated from Russian). Chayes, D., Kurras, G., Edwards, M., Anderson, R. & Coakley, B. J. 1998. Swath mapping the Arctic Ocean from US Navy submarines; Installation and performance analysis of SCAMP operation during SCICEX 1998. EOS, Transactions, American Geophysical Union, 79, F845. Chen, Y. M. 1982. Underwater acoustic ambient noise in the Arctic. Master thesis, Department of Ocean Engineering, Massachusetts Institute of Technology. Childers, V., McAdoo, D. C., Brozena, J. & Lon, S. 2001. New gravity data in the Arctic Occean: comparison of airborne and ERS gravity. Journal of Geophysical Research, 106, 8871–8886. Christensen, N. I. & Mooney, W. D. 1995. Seismic velocity structure and composition of the continental crust: a global view. Journal of Geophysical Research, 100, 9761–9788. Cochran, J. R., Edwards, M. H. & Coakley, B. J. 2006. Morphology and structure of the Lomonosov Ridge, Arctic Ocean. Geochemistry, Geophysics, Geosystems, 7, Q05019; doi: 10.1029/2005GC001114. Cochran, J. R., Kurras, G. J., Edwards, M. H. & Coakley, B. J. 2003. The Gakkel Ridge; bathymetry, gravity anomalies, and crustal accretion at extremely slow spreading rates. Journal of Geophysical Research, B, 108, 2116; doi: 10.1029/2002JB001830, EPM 12-1EPM 12-21. Coles, R. L. & Taylor, P. T. 1990. Magnetic anomalies. In: Grantz, A., Johnson, G. L. & Sweeeney, J. (eds) The Arctic Region. The Geology of North America, L. Geological Society of America, Boulder, CO, 119 –132. Cook, F. A., Coflin, K. C., Lane, L. S., Dietrich, J. R. & Dixon, J. 1987. Structure of the southeast margin of the Beaufort-Mackenzie Basin, Arctic Canada, from Crustal-Reflection Data. Geology, 15, 931 – 935. Crane, K., Eldholm, O., Myhre, A. M. & Sundvor, E. 1982. Thermal implication for the evaluation of the Spitsbergen Transform Fault. Tectonophysics, 89, 1 –32. Crary, A. A. 1954. Bathymetric chart of the Arctic Ocean along the route of T3, April 1952 to October 1953. Geological Society of American Bulletin, 65, 709 –712. Crary, A. P. & Goldstein, N. 1957. Geophysical studies in the Arctic Ocean. Deep-Sea Research, 4, 185 – 201. DeMets, C., Gordon, R. G., Argus, D. F. & Stein, S. 1990. Current plate motions. Geophysical Journal International, 101, 425 – 478.

699

Demenitskaya, R. M. & Hunkins, K. L. 1970. Shape and structure of the Arctic Ocean. In: Maxwell, A. E. (ed.) The Sea, 4, Part II. Wiley-Interscience, New York, 223 –249. Dietrich, J. R., Coflin, K. C., Lane, L. S., Dixon, J. & Cook, F. A. 1989. Interpretation of deep seismic reflection data, Beaufort Sea, Arctic Canada. Geological Survey of Canada, Ottawa, Open File Reports, 2106. Dietz, R. S. & Shumway, G. 1961. Arctic Basin Geomorphology. Geological Society of American Bulletin, 72, 1319– 1330. Doronin, Yu. P. 1970. Thermal Interaction of the Atmosphere and the Hydrosphere in the Arctic. Israel Program for Scientific Translations, Jerusalem. Double, M. J., Mercer, D. J. L & Meldrum, D. 2006. Wave measurements on sea ice: developments in instrumentation. Annals of Glaciology, 44, 108 –112. Dove, D. et al. 2009. Bathymetry, controlled source seismic, and gravity observations of the mendeleev ridge; implications for ridge structure, origin, and regional tectonics. Geophysical Journal International, submitted. Q11 Drachev, S. S., Johnson, G. L., Laxon, S. W., McAdoo, D. C. & Kassens, H. 1999. Main structural elements of Eastern Russian Arctic continental margin derived from satellite gravity and multichannel seismic reflection data. In: Kassens, H., Bauch, H. A. et al. (eds) Land–Ocean Systems in the Siberian Arctic: Dynamics and History. Springer, Berlin, 667 – 692. Dyer, I. 1984. The song of sea ice and other Arctic Ocean melodies. In: Dyer, I. & Chryssostomidis, C. (eds) Arctic Technology and Policy, Proceedings of the Second Annual MIT Sea Grant College Program Lecture and Seminar. Hemisphere, Washington, DC, 11– 37. Edwards, M. H. & Coakley, B. J. 2003. SCICEX investigations of the Arctic Ocean System. Chemie der Erde, 63, 281 –328. Engen, Ø., Faleide, J. I. & Dyreng, T. K. 2008. Opening of the fram strait gateway: a review of plate tectonic constraints. Tectonophys, 450, 51– 69. Q12 Engen, Ø., Gjengedal, J. A., Faleide, J. I., Kristoffersen, Y. & Eldholm, O. 2009. Seismic stratigraphy and basement structure of the Nansen Basin, Arctic Ocean. Geophysical Journal International, 176, 805 –821; doi: 10.1111/j.1365-246X.2008.04028.x. Feden, R. H., Vogt, P. R. & Fleming, H. 1979. Magnetic and bathymetric evidence for the ‘Yermak Hot Spot’ northwest of Svalbard in the Arctic Basin. Earth & Planetary, Science Letters, 44, 18– 38. Forsyth, D. A., Asudeh, I., Green, A. G. & Jackson, H. R. 1986. Crustal structure of the northern Alpha Ridge beneath the Arctic Ocean. Nature, 322, 349– 352. Franke, D., Hinz, K. & Oncken, O. 2001. The laptev sea rift. Marine Geology, 18, 1083– 1127. Fuetterer, D. K. 1992. Arctic ’91: the Expedition ARK-VIII/3 of RV ‘Polarstern’ in 1991. Berichte zur Polarforschung, 107, 267. Gaina, C., Roest, W. & Mu¨ller, R. D. 2002. Late cretaceous-cenozoic deformation of northeast Asia. Earth and Planetary Science Letters, 197, 273 –286. Geissler, W. H. & Jokat, W. 2004. A geopgysical study of the northern Svalbard continental margin. Geophysical Journal International, 158, 50– 66. Glebovsky, V. Yu., Likhachev, A. A., Kristoffersen, Y., Engen, Ø., Faleide, J. I. & Brekke, H. 2006. Sediment thickness estimations from magnetic data in the Nansen Basin. In: Scott, R. A. & Thurston, D. K. (eds) Proceedings of the Fourth International Conference on Arctic Margins, Dartmouth, Nova Scotia, 30 September to 3 October 2003. US Department of the Interior, US Minerals Management Service, Anchorage, AK, OCS Studies, MMS 2006-003, 157– 164. Glebovskiy, V. Yu., Minakov, A. N., Likhachev, A. A. & Merkur’ev, C. A. 2004. New summary map of Eurasian subbasin magnetic anomalies – a basis for geochronologic analysis of its formation history. In: Avetisov, G. P., Basov, V., Piskarev, A. L., Progebitsky, Yu. & Trukhalev, A. (eds) Geological – geophysical characteristics of the lithosphere in the Arctic Region. Proceedings of Ministry of Natural Resources of the Russian Federation, Russian Research Institute for Geology and Mineral Resources of

700

1036 1037 1038 1039 1040 1041 1042 1043 1044 1045 1046 1047 1048 1049 1050 1051 1052 1053 1054 1055 1056 1057 1058 1059 1060 1061 1062 1063 1064 1065 1066 1067 1068 1069 1070 1071 1072 1073 1074 1075 1076 1077 1078 1079 1080 1081 1082 1083 1084 1085 1086 1087 1088 1089 1090 1091 1092 1093 1094 1095 1096 1097 1098 1099 1100 1101 1102 1103 1104

Y. KRISTOFFERSEN

the Ocean. VNIIOkeangeologia, St Petersburg, 5, 62 –83 (translated from Russian). Gramberg, I. S., Kiselev, Yu. G. & Konovalov, V. V. 1991. Seismic studies from drift stations ‘North Pole’. Soviet Geology, N3, 45 – 54. Grantz, A., Clark, D. L. et al. 1998. Phanerozoic stratigraphy of Northwind Ridge, magnetic anomalies in the Canada Basin, and the geometry and timing of rifting in the Amerasia basin, Arctic Ocean. Geological Society of American Bulletin, 110, 801 –820. Grantz, A., Hart, P. E. & Childers, V. A. 2011. Geology and tectonic development of the Amerasia and Canada Basins, Arctic Occean. In: Spencer, A. M., Embry, A. F., Gautier, D. L., Stoupakova, A. V. & Sørensen, K. (eds) Arctic Petroleum Geology. Geological Society, London, Memoirs, 35, 773 –802. Grantz, A., Hart, P. E. et al. 1993. Cruise report and preliminary results U.S. geological Survey Cruise P1-93-AR Northwind Ridge and Canada Basin, Arctic Ocean aboard USCGC Polar Star, August 16– September 15, 1993. US Geological Survey, Reston, VA, Open-File Report, 93-389. Grantz, A., May, S. D. & Hart, P. E. 1990. Geology of the Arctic continental margin of Alaska. In: Grantz, A., Johnson, G. L. & Sweeney, J. F. (eds) The Arctic Ocean Region. The Geology of North America, L. Geological Society of America, Boulder, CO, 257– 289. Grantz, A., May, S. D. & Hart, P. E. 1994. Geology of the Arctic continental margin of Alaska. In: Plafker, G. & Berg, H. C. (eds) The Geology of Alaska. The Geology of North America, G-1. Geological Society of America, Boulder, CO, 17 –48. Grantz, A., Mullen, M. W., Philips, R. L. & Hart, P. E. 1992. Significance of extension, convergence, and transform faulting in the Chukchi Borderland for the Cenozoic tectonic development of the Arctic Basin. Geological Association of Canada Annual Meeting, 17, A42, Geological Association of Canada, St John’s, Newfoundland. Green, A. R., Kaplan, A. A. & Vierbuchen, R. C. 1986. Circum-Arctic Petroleum Potential. In: Future Petroleum Provinces of the World. American Association of Petroleum Geologists, Tulsa, OK, Memoirs, 40, 101– 129. Grikurov, G. E., Kaminsky, V. D. & Poselov, V. A. (eds) 2003. Morphology and geological nature of deep seabed and submarine elevations in the Arctic Basin: controversial scientific issues in context of UNCLOS/Article 76. Summary of Presentations International Conference, St. Petersburg, 30 June to 4 July 2003. Hall, J. K. 1973. Geophysical evidence for ancient sea-floor spreading from Alpha Cordillera and Mendeleyev Ridge. In: Arctic Geology. American Association of Petroleum Geologists, Tulsa, OK, Memoirs, 19, 542– 561. Hall, J. K. 1979. Sediment waves and other evidence of paleo-bottom currents at two locations in the deep Arctic Ocean. Sedimentary Geology, 23, 269 –299. Hall, J. K. 1990. Chukchi Borderland. In: Grantz, A., Johnson, G. L. & Sweeney, J. F. (eds) The Arctic Ocean Region. The Geology of North America, L. Geological Society of America, Boulder, CO, 337 – 350. Hall, J. K. 2008. Autonomous Arctic drifting Seismic buoys. HydroInternational, 12, December. Hall, J. K. & Kristoffersen, Y. 2009. ftputeba˚tR/H Sabvabaa – a research hovercraft for marine geophysical work in the most inaccessible area of the Arctic Ocean. The Leading Edge, August 2009, 932 – 935. Harris, R. A. 1904. Some indications of land in the vicinity of the North Pole. National Geographic, 15, 255 –261. Hunkins, K. 1960. Seismic studies of sea ice. Journal of Geophysical Research, 65, 3459– 3472. Hunkins, K. 1967. Initial oscillations of Fletcher’s Ice island (T-3). Journal of Geophysical Research, 72, 1165–1174. Hunkins, K. L. & Tiemann, W. 1977. Geophysical data summary for Fletcher’s Ice Island (T-3): May 1962– October 1974. Lamont – Doherty Survey of the World Ocean, Talwani, M. (ed.) LDGO Technical Report no. CU-1-77, May 1977, ONR Research Contract N00014-75-C-0505. Hunkins, K., Herron, E., Kutschale, H. W. & Peter, G. 1962. Geophysical studies of the Chukchi Cap, Arctic Ocean. Journal of Geophysical Research, 67, 235 – 247.

Jackson, H. R., Johnson, G. L., Sundvor, E. & Myhre, A. M. 1984. The Yermak Plateau: formed at a triple junction. Journal of Geophysical Research, 89, 3223– 3232. Jakobsson, M., Marcussen, C. LOMROG SCIENTIFIC PARTY 2008. Lomonosov Ridge off Greenland (LOMROG) 2007 Cruise Report. Geology Survey of Greenland, Copenhagen, Special Publications. Jakobsson, M., Mcnab, R. et al. 2008. An improved bathymetric potrayal of the Arctic Ocean: implications for ocean modeling and geological, geophysical and oceanographic analyses. Geophysical Research Letters; doi: 10.1029/2008g/033520. Q13 Jeffries, M. O. 1992. Arctic ice shelves and ice island: origin, growth and disintegration, physical characteristics, structuralvstratigraphic variability and dynamics. Reviews of Geophysics, 30, 245 –267. Johannessen, O. M., Shalina, E. V. & Miles, M. W. 1999. Satellite evidence for an Arctic sea ice cover in transformation. Science, 286, 1937– 1939. ˚ ., Gee, D. G., Larionov, A. N., Otha, Y. & Tebenkov, Johansson, A A. M. 2005. Grenvillian and Caledonian evolution of eastern Svalbard – a tale of two orogenies. Terra Nova, 17, 317 – 325. Johnson, G. L. 1983. The Fram Expeditions, Arctic Ocean Studies from Floating Ice, 1979–1982. Polar Record, 21, 583 –589. Jokat, W. 2003. Seismic investigations along the western sector of the Alpha Ridge, Central Arctic Ocean. Geophysical Journal International, 152, 185 –201. Jokat, W. 2005. The sedimentary structure of the Lomonosov Ridge between 88 N and 80 N. Geophysical Journal International, 163, 698 – 726. Jokat, W. & Micksch, U. 2004. Sedimentary structure of the Nansen and Amundsen basins, Arctic Ocean, Geophys. Research Letters, 31, L02603; doi: 10.1029/2003GL018352. Jokat, W., Buravtsev, V. Yu. & Miller, H. 1994. Marine seismic profiling in ice covered regions. Polarforschung, 64, 9 –17. Jokat, W., Geissler, W. & Voss, M. 2008. Basement structure of the north-western Yermak Plateau. Geophysical Research Letters, 35L05309, 1– 6. Jokat, W., Ritzmann, O., Schmidt-Aursch, M. C., Drachev, S., Gauger, S. & Snow, J. 2003. Geophysical evidence for reduced melt production on the Arctic ultraslow Gakkel mid-ocean ridge. Nature, 423, 962 – 965. Jokat, W., Uenzelmann-Neben, G., Kristoffersen, Y. & Rasmussen, T. 1992. Lomonosov Ridge – a double-sided continental margin. Geology, 20, 887 –890. Jokat, W., Weigelt, E., Kristoffersen, Y., Rasmussen, T. & Scho¨ne, T. 1995. New insights into the evolution of the Lomonosov Ridge and the Eurasia Basin. Geophysical Journal International, 122, 378 – 392. Judge, A. S. 1980. Heat flow measurements in the vicinity of the North Pole (abstr.). EOS Transactions, American Geophysical Union, 61, 277. Karasik, A. M. 1968. Magnetic anomalies of the Gakkel Ridge and the origin of the Eurasian sub-basin of the Arctic Ocean. Geophysical Survey Methods Arctic, 5, 8 – 19. Karasik, A. M. 1974. The Eurasia basin of the Arctic Ocean from the point of view of plate tectonics. In: Problems of Geology of Polar areas of the Earth. NIIGA, Leningrad, 23 – 31 (translated from Russian). Kenyon, S., Forsberg, R. & Coakley, B. 2008. New gravity field for the Arctic. EOS Transactions American Geophysical Union, 89; doi: 10.1029/2008EO320002. King, E. R., Zietz, I. & Alldredge, L. R. 1964. Genesis of the Arctic Ocean basin. Science, 144, 1551–1557. Klemperer, S. L., Miller, E. L., Grantz, A., Scholl, D. & BERINGCHUKCHI WORKING GROUP 2002. Crustal structure of the Bering and Chukchi shelves: deep seismic reflection profiles across the North American continent between Alaska and Russia. In: Miller, E. L., Grantz, A. & Klemperer, S. L. (eds) Tectonic Evolution of the Bering Shelf-Chukchi Sea-Arctic Margin and Adjacent Landmasses. Geological Society of America, Boulder, CO, Special Papers, 360, 1– 24. Koenig, L. S., Greenaway, K. R., Dunbar, M. & Hattersley-Smith, G. 1952. Arctic Ice Islands. Arctic, 5, 67 –103.


1105 Kovacs, L. C., Glebovsky, V. Yu., Maschenkov, S. P. & Brozena, J. 2002. New map and grid of completed magnetic anomalies from 1106 the Arctic Ocean. EOS, Transactions American Geophysical Union, 1107 Fall Meeting Supplement, 83, 1330; World Wide Web Address: 1108 http://www.geus.dk/program-areas/nature 1109 1110 Kristoffersen, Y. 1997a. Seismic Reflection Surveys during Arctic Ocean-96. Yearbook 1995/96. Swedish Polar Research Secretariat, 1111 75 – 77. 1112 1113 Kristoffersen, Y. 1997b. A Pilot Project for Shallow Drilling in the Arctic Ocean. Yearbook 1995/96. Swedish Polar Research Sec1114 retariat, 72 – 74. 1115 Kristoffersen, Y. & Husebye, E. S. 1985. Multichannel seismic reflec1116 tion measurements in the Eurasia Basin, Arctic Ocean, from U.S. ice 1117 station FRAM-IV. Tectonophysics, 114, 03 –117. 1118 Kristoffersen, Y. & Mikkelsen, N. (eds) 2004. Scientific drilling in the 1119 Arctic Ocean and the site survey challenge: Tectonic, paleoceano1120 graphic and climatic evolution of the Polar Basin. In: JEODI Work1121 shop, Copenhagen, January 2003. Geology Survey of Denmark and 1122 Greenland, Copenhagen, Special Publications. http://www.geus. 1123 dk/program-areas/nature-environment/international/reports/ 1124 Kristoffersen, Y. & Mikkelsen, N. 2006. On sediment deposition and nature of the plate boundary at the junction between the submarine 1125 Lomonosov Ridge, Arctic Ocean and the continental margin of 1126 Arctic Canada/North Greenland. Marine Geology, 22, 265 –278. 1127 1128 Kristoffersen, Y., Backman, A. & Brass, G. 1992. Logistics of scientific drilling in the Arctic Ocean: station-keeping test opens new 1129 perspectives. The Nansen Icebreaker, 3, 1. 1130 1131 Kristoffersen, Y., Coakley, B., Jokat, W., Edwards, M., Brekke, H. & Gjengedal, J. 2004. Seabed erosion on the Lomonosov Ridge, 1132 central Arctic Ocean: a tale of deep draft icebergs in the Eurasia 1133 basin and the influence of Atlantic water inflow on iceberg motion? 1134 Paleoceanography, 19, PA3006; doi: 10.1029/2003PA000985. 1135 Kristoffersen, Y., Husebye, E. S., Bungum, H. & Gregersen, S. 1982. 1136 Seismic investigations of the Nansen Ridge during the FRAM-I 1137 experiment. Tectonophysics, 82, 57 –68. 1138 Kutschale, H. 1969. Arctic Hydroacoustics. Arctic, 22, 246– 264. 1139 Lachenbruch, A. H. & Marshall, B. V. 1966. Heat flow through 1140 the Arctic Ocean floor: the Canada Basin-Alpha Rise boundary. 1141 Journal of Geophysical Research, 71, 1223–1248. 1142 Lachenbruch, A. H. & Marshall, B. V. 1969. Heat flow in the Arctic. 1143 Arctic, 22, 300 – 311. 1144 Lane, L. S. 2002. Tectonic ecolution of the Canadian Beaufort SeaMackenzie Delta region: a brief review. Canadian Society of 1145 Exploration Geophysicists Recorder, February, 49 – 55. 1146 1147 Langinen, A. E., Lebedeva-Ivanova, N. N., Gee, D. G. & Zamansky, Yu. Ya. 2008. Correlations between the Lomonosov Ridge, Marvin 1148 Spur, and adjacent basins of the Arctic Ocean based on seismic 1149 data. Tectonophys, xx, xx– xx. Q14 1150 1151 Laxon, S. W. & McAdoo, D. C. 1998. Satellites provide insights into polar geophysics. EOS Transactions, American Geophysical Union, 1152 79, 69, 72 – 73. 1153 Lebedeva-Ivanova, N. N., Zamansky, Yu. Ya., Langinen, A. E. & 1154 Sergeyev, M. B. 2006. A seismic model for the Earth’s crust along 1155 the “East-Siberian Continental Margin– Podvodnikov Basin-Arlis 1156 Rise” Geotraverse (Arctic Ocean). In: Scott, R. & Thurston, D. 1157 (eds) Proceedings of the Fourth International Conference on Arctic 1158 Margins, Dartmouth, Nova Scotia, 30 September to 3 October 1159 2003, 132 –135. 1160 Levin, F. K. 1983. The effects of streamer feathering on stacking. Geo1161 physics, 48, 1165–1171. 1162 Litvinova, T. & Glebovsky, V. 2008. New map and grid of compiled 1163 magnetic nomalies for the Arctic Icean an adjacent continental part of the Russian Federation. Geophysical Research Abstracts, 10, 1164 EGU2008-A-12245. Q15 1165 1166 Lubimova, E. A., Alexandrov, A. L. & Duchkov, A. D. 1973. Methods of Study of the Heat Flow Through the Bottom of the Ocean. 1167 Nauka, Moscow. 1168 1169 Lubimova, E. A., Nikitina, V. N. & Tomara, G. A. 1976. Thermal Fields of the Inland Marginal Seas of the USSR. Nauka, Moscow. 1170 1171 Makris, N. C. & Dyer, I. 1986. Environmental correlates of pack ice noise. Journal of Acoustical Society of America, 79, 1434–1440. 1172 1173

701

Marcussen, C. 2008. Lomonosov ridge off Greenland (LOMROG) 2007: Danish continental shelf project. Yearbook 2007 Swedish Polar Secretariat, 137 –141. Maschenkov, S. P., Glebovsky, V. Yu. & Zayoncheck, A. V. 1999. New digital compilation of Russian aeromagnetic and gravity data over the North Eurasian shelf. Polarforschung, 69, 35 –39. Maykut, G. A. & Untersteiner, N. 1971. Some results from a time dependent, thermodynamic model of sea ice. Journal of Geophysical Research, 76, 1550–1575. Meyer, O., Peddie, D. & Kristoffersen, Y. 2008. Ny bøye vil gi mer seismiske data fra Polhavet. GEO, 8, 53. Michael, P. J., Langmuir, C. H. et al. 2003. Magmatic and amagmatic seafloor generation at the ultraslow-spreading Gakkel Ridge, Arctic Ocean. Nature, 423, 956– 961. Mickey, M. B. & Haga, A. P. 1987. Jurassic-Neocomian biostratigraphy, North Slope, Alaska. In: Tailleur, I. L. & Weimer, P. (eds) Alaskan North Slope Geology, Vol. 1. Pacific Section Society of Economic Paleontologist and Mineralogists and Alaskan Geological Survey, Bakersfield, CA, 397– 404. Mickey, M. B., Byrnes, A. P. & Haga, H. 2002. Biostratigraphic evidence for the prerift position of the North Slope, Alaska, and Arctic Islands, Canada, and Sinemurian incipient rifting of the Canada Basin. In: Miller, E. L., Grantz, A. & Klemperer, S. L. (eds) Tectonic Evolution of the Bering Shelf –Chukchi Sea – Arctic Margin and Adjacent Landmasses. Geological Society of America, Boulder, CO, Special Papers, 360, 67– 75. Moran, K. et al. 2006. The Cenozoic palaeo environment of the Arctic Ocean. Nature, 441; doi: 10.1038nature. Q16 Nansen, F. 1904. The bathymetrical features of the North Polar seas, with a discussion of the continental shelves and previous oscillations of the shore-line. In: Nansen, F. (ed.) The Norwegian North Polar Expedition 1893–1896. Scientific Results. Christiania (Jacob Dybwad), Vol. IV, part 13. Q17 Ostenso, N. A. & Wold, R. J. 1977. A seismic and gravity profile across the Arctic Ocean basin. Tectonophys, 37, 1– 24. Ostrekin, M. E. 1954. Latest investigations of Central Arctic. Priroda, 12, 3– 12 (in Russian, translation by E.R. Hope, September 1955). Plouff, D., Keller, G. V., Frischknecht, F. C. & Wahl, R. R. 1961. Geophysical Studies on IGY Drifting Station Bravo, T-3, 1958 to 1959. In: Raasch, G. O. (ed.) Geology of the Arctic. Proceedings of the First International Symposium on Arctic Geology. Calgary, Alberta, 11– 13 January 1960, 1, 709 – 716. Popelar, J. & Kouba, J. 1983. Satellite doppler determination of differential sea ice motion in the vicinity of the North pole. Marine Geodesy, 7, 171 –202. Proshutinsky, A. Y. & Johnson, M. A. 1997. Two circulation regimes of the wind driven circulation. Journal of Geophysical Research, 102, 12,394– 12,514. Rassokho, A. I., Senchura, L. I., Demenitskaya, R. M., Karasik, A. M., Kiselev, Yu. G. & Timoshenko, N. K. 1966. Submarine MidArctic Ridge and its place in the Arctic Ocean ridge system. Doklady Akademii Nauk S.S.S.R., 172, 659 – 662. Rigor, I. G., Wallace, J. M. & Colony, R. L. 2002. Response of sea ice to the Arctic Oscillation. Journal of Climate, 15, 2648–2663. Rothrock, D. A. & Thorndike, A. S. 1980. Geometric properties of the underside of sea ice. Journal of Geophysical Research, 85, 3955–3963. Rothrock, D. A., Yu, Y. & Maykut, G. A. 1999. Thinning of the Arctic sea-ice cover. Geophysical Research Letters, 26, 3469– 3472. Saltus, R. W. & Bird, K. J. 2003. Digital depth horizon compilations of the Alaskan North Slope and adjacent arctic regions. US Geological Survey, Reston, VA, Open-File Reports, 03-230. Schwindlein, V., Mu¨ller, C. & Jokat, W. 2007. Microseismicity of the ultraslow-spreading Gakkel Ridge, Arctic Ocean: a pilot study. Geophysical Journal International 169, 100 –112. Sekretov, S. B. 2001. Northwestern margin of the East Siberian Sea, Russian Arctic: seismic stratigraphy, structure of the sedimentary cover and some remarks on the tectonic history. Tectonophys, 339, 353– 383.

702

1174 1175 1176 1177 1178 1179 1180 1181 1182 1183 1184 1185 1186 1187 1188 1189 1190 1191 1192 1193 1194 1195 1196 1197 1198 1199 1200 1201 1202 1203 1204 1205 1206 1207 1208 1209 1210 1211 1212 1213 1214 1215 1216 1217 1218 1219 1220 1221 1222 1223 1224 1225 1226 1227 1228 1229 1230 1231 1232 1233 1234 1235 1236 1237 1238 1239 1240 1241 1242

Y. KRISTOFFERSEN

Sekretov, S. B. 2002. Structure and tectonic evolution of the Southern Eurasia Basin, Arctic Ocean. Tectonophysics, 351, 192– 243. Serreze, M., Box, J., Barry, R. & Walsh, J. 1993. Characteristics of Arctic synoptic activity. Meterology and Atmospheric Physics, 51, 147 – 167. Shepard, G. W. 1979. Arctic Ocean Ambient Noise. Master thesis, Department of Ocean Engineering, Massachusetts Institute of Technology. Sobczak, L., Mayr, U. & Sweeney, J. F. 1986. Crustal section across the polar continent-ocean transition in Canada. Canadian Journal of Earth Sciences, 23, 608– 621. Sohn, R. et al. 2008. Explosive volcanism on the ultraslow-spreading Gakkel ridge, Arctic Ocean. Nature, 453, 1236– 1238. Sorokin, M. Yu., Zamansky, Yu. Ya., Langinen, A. Ye., Jackson, H. R. & Macnab, R. 1999. Crustal structure of the Makarov Basin, Arctic Ocean determined by seismic refraction. Earth & Planetary Science Letters, 168, 187 – 199. Stephenson, R. A., Coffin, K. C., Lane, L. S. & Dietrich, J. R. 1994. Crustal structure and tectonics of the southeastern Beaufort Sea continental margin. Tectonics, 13, 389 – 400. Sverdrup, H. U. 1931. Hvorledes og Hvorfor med Nautilus. Gyldendal Norsk Forlag, Oslo. Sweeney, J. F., Sobczak, L. & Forsyth, D. A. 1990. The continental margin northwest of the Queen Elizabeth Islands. In: Grantz, A., Johnson, G. L. & Sweeney, J. F. (eds) The Arctic Ocean Region. The Geology of North America, L. Geological Society of America, Boulder, CO, 227– 238. Taylor, A. E., Judge, A. & Allen, V. S. 1986. Terrestrial heat flow from Project CESAR; Alpha Ridge, Arctic Ocean Basin. Journal of Geodynamics, 6, 137 –176. Taylor, P. T., Kovacs, L. C., Vogt, P. R. & Johnson, G. L. 1981. Detailed aeromagnetic investigation of the Arctic Basin, 2. Journal of Geophysical Research, 86, 6323–6333. Thiede, J. 1988. Scientific cruise report of Arctic Expedition ARK IV/3. Berichte zur Polarforschung, 43. Tieman, W., Ardai, J., Allen, B., Manely, T. O., Kristoffersen, Y. & Hunkins, K. 1982. Geophysical data from drifting stations FRAM IV and TRISTEN. Lamont-Doherty Geological Observatory of Columbia University, New York, Technical Report, LDGO-82-3. Thompson, D. W. J. & Wallace, J. M. 1998. The Arctic Oscillation signature in the wintertime geopotential height and temperature fields. Geophysical Research Letters, 25, 1297– 1300. Tucker, III, W. 1989. An overview of the physical properties of sea ice. In: Offshore Workshop on Ice Properties, Associate Committee on Geotechnical Research, National Research Council Canada (June 21 –11), 1988, St Johns, Newfoundland, Technical memorandum no. 44, NRCC 30358. Q18 Untersteiner, N. 1990. Structure and dynamics of the Arctic Ocean ice cover. In: Grantz, A., Johnson, G. L. & Sweeeney, J. (eds) The

Arctic Region. The Geology of North America, L. Geological Society of America, Boulder, CO, 37 –62. Untersteiner, N. & Thorndike, A. S. 1982. Arctic data buoy program. Polar Record, 21, 127 –135. Untersteiner, N., Hunkins, K. L. & Buck, B. M. 1976. Arctic science: current knowledge and future thrusts. In: Science, Technology and the Modern Navy. Thirtieth Anniversary 1946– 1976. Office of Naval Research, Arlington, VA, 0NR-37, 278– 297. Vancoppenolle, M., Fichfet, T. & Bitz, C. M. 2006. Modeling the salinity profile of undeformed Arctic sea ice. Geophysical Research Letters, 33, L21501; doi: 10.1029/2006GL028342. Verhoef, J., Roest, W. R., Macnab, R., Arkani-Hamed, J. & MEMBERS OF THE PROJECT TEAM 1996. Magnetic anomalies of the Arctic and North Atlantic Oceans and adjacent land areas. Geological Survey of Canada, Ottawa, Open File Reports, 3125, Parts A and B. Vogt, P. R., Taylor, P. T., Kovacs, L. C. & Johnson, G. L. 1979. Detailed investigations of the Arctic Basin. Journal of Geophysics Research, 84, 1071– 1089. Vogt, P. R., Taylor, P. T., Kovacs, L. C. & Johnson, G. L. 1982. The Canada Basin: aeromagnetic constraints on structure and evolution. Tectonophysics, 89, 295 –336. Wadhams, P. 1994. Sea Ice Thickness Changes and Their Relation to Climate. In: Johannessen, O. M., Munch, R. D. & Overland, J. (eds) The Polar Oceans and Their Role in Shaping the Global Environment. American Geophysics Union, Washington, DC, Geophysics Monographs, 85, 337 – 363. Watanabe, E. & Hasumi, H. 2005. The Arctic sea ice response to wind stress variations. Journal of Geophysics Research, 110, C11007; doi: 10.1029/2004JC002678. Weber, J. R. 1979. The lomonosov ridge experiment: ‘Lorex 79’. EOS, Transactions, American Geophysical Union, 60, 715 –721. Weber, J. R. 1983. Maps of the arctic basin sea floor: a history of bathymetry and its interpretation. Arctic, 36, 121 – 142. Weeks, W. F., Kovacs, A. & Hibler, III, W. D. 1971. Pressure ridge characteristics in the Arctic coastal environment. In: Wetteland, S. & Bruun, P. (eds) Proceedings First International Conference on Port Ocean Engineering under Arctic Conditions, 1, 152– 183. Weigelt, E. & Jokat, W. 2001. Pecularities of roughness and thickness of oceanic crust in the Eurasian Basin, Arctic Ocean. Geophysics Journal International, 145, 505 – 516. Wessel, P. & Smith, W. 1991. Free software helps map and display data. EOS, 72, 441, 445 –446. Wilson, J. T. 1963. Hypothesis of the Earth’s behaviour. Nature, 198, 925 – 929. Worthington, L. V. 1953. Oceanographic results of project ski jump I and ski jump II in the polar sea, 1951–1952. Transactions, American Geophysical Union, 34, 543 –551.

Author Query Form

Author Query Form

Suggest Documents