The Deep Waters of the Eurasian Basin, Arctic Ocean: Geothermal Heat Flow, Mixing and Renewal

ARTICLE IN PRESS

Deep-Sea Research I 53 (2006) 1253–1271 www.elsevier.com/locate/dsr

The deep waters of the Eurasian Basin, Arctic Ocean: Geothermal heat ﬂow, mixing and renewal$

Go¨ran Bjo¨rka,Ã, Peter Winsorb

aDepartment of Oceanography, Earth Sciences Centre, Go¨teborg University, P.O. Box 460, 405 30 Go¨teborg, Sweden bPhysical Oceanography Department, Woods Hole Oceanographic Institution, Woods Hole, MA 02543, USA

Available online 12 July 2006

Abstract

Hydrographic observations from four separate expeditions to the Eurasian Basin of the Arctic Ocean between 1991 and 2001 show a 300–700 m thick homogenous bottom layer. The layer is characterized by slightly warmer temperature compared to ambient, overlying water masses, with a mean layer thickness of 5007100 m and a temperature surplus of 7.072 103 1C. The layer is present in the deep central parts of the Nansen and Amundsen Basins away from continental slopes and ocean ridges and is spatially coherent across the interior parts of the deep basins. Here we show that the layer is most likely formed by convection induced by geothermal heat supplied from Earth’s interior. Data from 1991 to 1996 indicate that the layer was in a quasi steady state where the geothermal heat supply was balanced by heat exchange with a colder boundary. After 1996 there is evidence of a reformation of the layer in the Amundsen Basin after a water exchange. Simple numerical calculations show that it is possible to generate a layer similar to the one observed in 2001 in 4–5 years, starting from initial proﬁles with no warm homogeneous bottom layer. Limited hydrographic observations from 2001 indicate that the entire deep-water column in the Amundsen Basin is warmer compared to earlier years. We argue that this is due to a major deep-water renewal that occurred between 1996 and 2001. r 2006 Elsevier Ltd. All rights reserved.

Keywords: Deep water; Convection; Mixing; Geothermal heat; Heat ﬂow; Nansen Basin; Amundsen Basin; Lomonosov Ridge; Arctic Ocean

1. Introduction tions, quite different from the uniform mixing rate assumed in models. While we understand much Diapycnal mixing in the oceans is a prerequisite about the processes responsible for mixing, and for the renewal of the deep waters and is essential have a rough qualitative sense of their distribution, for upper ocean heat transfer and storage. Under- quantitative information on mixing is lacking for standing this critical process has become more most of the world oceans. This is especially true for urgent, as it is now realized that mixing processes the ice-covered and poorly sampled Arctic Ocean. have heterogeneous and complex spatial distribu- Observations of mixing in the upper ocean of the Arctic show that energy levels associated with $Woods Hole Oceanographic Institution contribution 11081. internal waves are roughly 1/3 of those found in ÃCorresponding author. Tel.: +46 31 7732958. mid- and low-latitude regions of the oceans E-mail address: [email protected] (G. Bjo¨rk). (D’Asaro and Morehead, 1991; Plueddemann,

1992). It has been argued that more subtle, double- respectively, divided by the Gakkel Ridge having a diffusive processes may control much of the interior broken structure and generally low altitudes. The mixing (e.g. Padman, 1994). However, processes only direct deep-water communication is with the responsible of abyssal mixing and ventilation in the Greenland-Iceland-Norwegian (GIN) seas through Arctic are still relatively unknown. Studies suggest Fram Strait with a sill depth of about 2500 m (see that the deep waters may be ventilated slowly by Fig. 1 for geographical information). Thus, the renewal of shelf water created by freezing and brine Arctic deep waters below 2500 m are isolated from rejection on the shelves (e.g. Aagaard et al., 1981; direct influence of the surrounding oceans. The Rudels et al., 2000; Winsor and Bjo¨rk, 2000)or estimated isolation ages for the Eurasian Basin and from influxes from the adjacent Norwegian Sea Canadian Basin bottom waters, based on 14C (Aagaard et al., 1985; Jones et al., 1995). measurements, are 250 and 450 years, respec- There are also topographic constraints that tively (Schlosser et al., 1997). determine the type of processes that may affect the The isolated nature of the deep waters suggests deep-water mixing. The Arctic Ocean is divided into that relatively weak fluxes may affect the deep-water two main basins: the Eurasian Basin and the properties over time. One possible flux is heat Canadian Basin, separated by the central Lomono- supplied to the deep ocean from the Earth’s interior, sov Ridge. The Eurasian Basin consists of two sub- the geothermal heat flow. Most observations of basins: the Nansen and Amundsen Basins with geothermal heat flux have focused on extreme fluxes typical maximum depth of about 4000 and 4500 m, at small point sources, such as hydrothermal vents,

Fig. 1. Map showing the positions of the CTD stations used in the paper. Squares (black) are from Oden’91 (1991), triangles (green) AOS94, circles (red) ARKXII (1996), and stars (blue) from AO-01 (2001). Filled markers indicate where a homogeneous bottom layer was found. Bathymetry is from the International Bathymetric Atlas of the Arctic Ocean (IBCAO). ARTICLE IN PRESS G. Bjo¨rk, P. Winsor / Deep-Sea Research I 53 (2006) 1253–1271 1255 but there is also a background heat flux associated Arctic: the Oden’91 cruise on I/B Oden (1991), the with the warm interior of the Earth. Stein and Stein AOS94 cruise on CCGS Louis S. St-Laurent (1994), (1992) estimated a global mean geothermal heat flux the ARK XII expedition of 1996 with Polarstern, of 87 mW m2 (1 mW ¼ 103 W), with a back- and the AO-01 expedition in 2001 with I/B Oden. ground level of about 50 mW m2 over deep abyssal Potential temperature Y (referenced to the surface plains. Mid-ocean ridges generally have larger unless otherwise noted) and potential density s were values 4200 mW m2 (Murton et al., 1999). computed from standard algorithms (UNESCO, Although the geothermal heat flux is relatively 1983). Fig. 1 shows a map of the Eurasian part of small compared to other fluxes, it has been shown in the Arctic Ocean with bathymetry and positions of several studies that it may have a significant effect the different hydrographic stations used here. on the large-scale circulation. Gustafsson (2002) used the mean global geothermal heat flux to estimate the production of potential energy and 2.2. Geothermal heat flow compared this with estimates of the total amount of energy available from tides and winds for deep Observations of geothermal heat flow are based water mixing. She found that the heating of the deep on measurements of the temperature gradient in the waters of the world’s oceans by geothermal heat sediment. The heat flow Q is determined from flow represents about 4% of the heating from Q ¼ kdT/dz, where k is the thermal conductivity, vertical diffusion, but that the actual work against and dT/dz is the temperature gradient in the buoyancy forces done by geothermal heating can be sediment, usually measured by a probe inserted 10% of the total work by turbulent diffusion. The into the bottom, e.g. a gravity corer. Most published geothermal heat flux can therefore be an important data from the Arctic region comes from the component of the global thermohaline circulation. Canadian Basin, and is available through the This was also shown in modeling studies by Adcroft Global Heat Flow Data Set stored at the US et al. (2001) and Scott et al. (2002). They used a National Geophysical Data Center (NGDC). A spatially uniform heat flux (50 mW m2) through general presentation of this extensive data set is the ocean floor to investigate its effect on the given in Pollack et al. (1993). meridional overturning circulation and found that Langseth et al. (1990) presented an overview of the imposed heat flux caused a perturbation of the geothermal heat flow measurements in the Arctic, meridional overturning cell on the order of several with most observations originating from the Cana- Sverdrups, connecting with an upper level circula- dian Basin. They found typical heat fluxes ranging tion at high latitudes where the imposed heat flow between 30 and 105 mW m2, with an overall mean altered the deep-water circulation in the oceans of 46 mW m2. Majorowicz and Embry (1998) significantly. estimated heat flows in the 35–90 mW m2 range, In this paper we use hydrographic measurements with a mean of 53712 mW m2 from 156 sites in the from icebreaker expeditions to the Arctic from 1991 Sverdrup Basin, Canadian Arctic. Earlier observa- to 2001, focusing on the deep-water properties of tions over the Alpha Ridge during the Cesar the Eurasian Basin, and in particular, the vertically experiment found an average heat flow of homogeneous bottom layer presumably caused by 60 mW m2 (Taylor et al., 1986). Recent measure- geothermal heat flow. Since many of the processes ments on the continental slope of the Laptev Sea that affect the observed layer in the Eurasian Basin revealed relatively high heat flow values between 85 are unknown, we investigate the dynamics of this and 117 mW m2 (Drachev et al., 2003). layer in an exploratory fashion as a starting point Observations from the central Eurasian Basin are for future investigations. relatively sparse. The data used in this paper come from yet unpublished material collected during the 2. Observations ARK VI, ARK XVI, and ARK XVII cruises (N. Kaul personal communication, 2002) and consist of 2.1. Hydrographic data 15 individual measurements transecting the central Amundsen Basin from the Gakkel Ridge to the The hydrographic observations consist of high Lomonosov Ridge. From these data we find a vertical resolution conductivity–temperature–depth relatively constant heat flow of 80 mW m2 over the (CTD) data from four icebreaker expeditions to the Amundsen Basin, which we use as a representative ARTICLE IN PRESS 1256 G. Bjo¨rk, P. Winsor / Deep-Sea Research I 53 (2006) 1253–1271 average value for the Nansen and Amundsen Basins shows an example of vertical profiles of Y and throughout this paper. salinity S from the AO-01 cruise (station 26). The profiles are from the Nansen Basin, located at 841N 2.3. Evidence of geothermal heating of bottom layers 331E (see Fig. 1). Panels 2c and 2d show the detailed structure of the lowest 1000 m. The potential We focus on the deepest 1000 m of the ocean, and temperature plotted in Fig. 2c shows that there is in particular the presence of a vertically homo- an apparent, vertically homogenous layer extending genous bottom layer. This layer was ‘‘discovered’’ from the bottom and up. The vertical extent of the while vertical CTD profiles during the AO-01 were layer, H, is about 550 m, and the potential processed. Examination of earlier data from the temperature difference DY compared to the over- Eurasian Basin revealed that such a layer was lying water, is 0.012 1C. The temperature differ- present on all four cruises investigated here. Fig. 2 ence is defined as the difference between the top of

1000 1000 (a) (b)

1500 1500

2000 2000

2500 2500

Pressure (dbar) 3000 3000

3500 3500

2∆S 4000 4000 -0.9 -0.7 -0.5 -0.3 34.91 34.93 34.95 Θ (°C) S 3000 3000 (c) (d)

3200 3200

3400 3400

3600 ∆Θ 3600 Pressure (dbar)

3800 H 3800

4000 4000

-0.943 -0.939 -0.935 -0.931 34.939 34.942 34.945 Θ (°C) S

Fig. 2. Vertical profiles of (a) potential temperature Y, (b) salinity S from the Nansen Basin as observed during the AO-01 expedition (station 26, 84144.00N, 35114.90E, 18 July, 2001). Panels c and d show the bottom 1000 m of the profiles with exaggerated scale. Panel c also shows the definition of the layer height H, and the potential temperature anomaly DY. The salinity deficit DS is defined (see panel b as the difference between the observed salinity in the mixed layer and the salinity that results if the profile above the mixed layer is extrapolated to the bottom and mixed over the distance H. ARTICLE IN PRESS G. Bjo¨rk, P. Winsor / Deep-Sea Research I 53 (2006) 1253–1271 1257 the homogenous layer and the temperature mini- heating) would not have any excess temperature mum located a couple of hundred meters above this since the temperature in the Eurasian Basin below layer. Values of H and DY for some selected 1000 m generally decreases with depth. Another stations are given in Table 1. The heat surplus in possible formation mechanism is that dense and the homogeneous layer suggests that it is formed by relatively warm gravity currents from the shelves geothermal heat supplied from the Earth’s interior, might reach the deepest parts of the basin and and the constant potential temperature and salinity generate a well mixed and warmer layer, but in through the layer indicate ongoing convection. The order to reach all the way to the bottom they need step-like structure immediately above the layer is negative buoyancy and must then be saltier than the seen at many stations (e.g. Fig. 2c) and is probably original bottom water. There is no sign of an caused by double diffusive processes. Such step enhanced salinity in the observed bottom layer. On features are also found in the isolated deep waters of the contrary there is a salinity deficit (DS in Fig. 2b) the Canada Basin (Timmermans et al., 2003). in the mixed layer that ranges between 0.1 103 The temperature excess in the well-mixed bottom and 4.0 103 for all stations with a well-mixed layer is by itself a strong indicator that this layer has bottom layer. It is also likely that inflows of new been formed by heating from below. A well-mixed bottom water occur intermittently with a long layer formed by only mechanical mixing (with no period in between when diffusion processes would

Table 1 Statistical properties of convective layers in Arctic bottom waters from four cruises

Cruise Station Position Basin Depth DY H

Lat. (1N) Long. (1E) N*/A* (m) (103 C) (m)

Oden’91 (1991) 8 83133.30 27137.60 N 4050 5.5 400 985104.30 42118.10 N 4000 3.5 500 16 87136.20 69144.70 A 4450 6.8 600 17 88100.30 85103.30 A 4400 8.0 625 18 88110.90 99107.70 A 4410 4.0 360 30 88159.30 8156.30 A 4370 6.0 470 31 88116.60 9120.40 A 4415 7.5 555 32 87129.70 11144.70 A 4377 8.0 375 33 86145.30 10107.60 A 4382 10.0 460 35 86109.60 05113.80 A 4435 10.0 500

AOS94 (1994) 35 89159.90 66159.50 A 4300 3.5 335 37 84114.70 35100.50 N 4000 8.0 450 38 83150.70 35141.40 N 4050 8.0 450 ARKXII (1996) 41 83130.00 96134.80 N 3616 10.0 390 43 84112.10 100132.00 N 3777 8.0 440 50 85110.10 109117.30 A 4450 5.5 675 51 85117.00 111135.20 A 4420 5.5 470 52 85124.30 113100.60 A 4450 6.1 690 54 85145.50 117128.20 A 4450 7.0 650 55 85152.70 121114.10 A 4440 6.3 640 56 86109.60 125148.80 A 4410 6.8 610 57 86109.10 125156.80 A 4370 6.0 610 AO-01 (2001) 26 84144.00 35114.90 N 4050 12.0 550 36 88126.00 109150.60 A 4425 5.0 405 54 88144.70 1140.00 A 4410 6.0 460 60 88115.10 8109.60 A 4425 6.0 525

Mean 7std 7.072.1 5027105

DY is the potential temperature difference between the convective layer and the overlying water mass, and H is the thickness of the convective layer. N denotes Nansen Basin and A denotes Amundsen Basin. See Fig. 1 for station locations and Fig. 2 and text for further details. ARTICLE IN PRESS 1258 G. Bjo¨rk, P. Winsor / Deep-Sea Research I 53 (2006) 1253–1271 re-stratify the bottom layer. Such a stratification Fig. 3 shows vertical profiles of potential tem- would contradict the finding of a well-mixed layer perature from the Oden’91 expedition along the during several consecutive expeditions. The pre- easternmost section, extending from the Barents Sea sence of a significant geothermal heat flux in the shelfbreak to the Lomonosov Ridge (stations 5–19 Eurasian Basin supports the idea that the layer is in Fig. 1). A homogenous layer is present in both formed by geothermal heat flux that heats up water the Nansen and Amundsen Basins. Note how the near the bottom, maintaining convective motions, layer is present in the main deep basins, but not in and keeping the bottom 300–700 m well mixed. areas with pronounced topographic variability, e.g. the Gakkel and Lomonosov Ridges. Fig. 4a shows a hydrographic section from the 2.4. Temporal and spatial features ARKXII (1996) cruise, located further to the east than the 1991 section. Again, we find a well- Data from the different icebreaker expeditions in developed bottom layer, especially visible in the the Eurasian Basin enable us to examine the spatial Amundsen Basin. The thickness of the layer is and temporal scales of the homogenous bottom 700 m and occupies the width of the interior basin. layer. Two expeditions, Oden’91 and AO-01, Fig. 4b shows a detailed plot of the ARKXII traversed the Amundsen Basin twice during each profiles across the Amundsen Basin (stations expedition (on the way to and from the North Pole), 52–57). making it possible to examine the spatial extent of Filled markers in Fig. 1 show the spatial the homogenous layer. distribution of all observations with a homogenous

Θ (°C) -0.94 -0.92 -0.9 -0.88 - 0.86 -0.84 -0.82 3000 8 10 11 12 13 14 16 17 18 19

3500 Pressure (dbar)

4000

NB GR AB LR 4500

Fig. 3. Vertical profiles of potential temperature, Y, from the Oden’91 expedition across the Nansen and Amundsen Basins. Station numbers are shown at the top of each profile. See Fig. 1 for locations. Bold numbers indicate stations with a well defined bottom mixed layer. Deep basins and ocean ridges along the section are also indicated. NS ¼ Nansen Basin; GR ¼ Gakkel Ridge; AB ¼ Amundsen Basin; LR ¼ Lomonosov Ridge. Successive profiles are offset by 0.01 1C. ARTICLE IN PRESS G. Bjo¨rk, P. Winsor / Deep-Sea Research I 53 (2006) 1253–1271 1259

Θ (°C) 0.95 0.9 0.85 0.8 3000 42 43 44 45 46 47 48 49 50 52 55 56 57 58 59 54 51 41

3500 Pressure (dbar)

4000

(a) NB GR AB LR 4500 Θ (°C)

0.85 0.84 0.83 0.82 0.81 0.8 3000 52 54 55 56 57

3500 Pressure (dbar)

4000

(b) 4500

Fig. 4. (a) Vertical profiles of potential temperature, Y, from the ARK XII (1996) expedition across the Nansen and Amundsen Basins. Station numbers are shown at the top of each profile. See Fig. 1 for locations. Bold numbers indicate stations with a well defined bottom mixed layer. Successive profiles are offset by 0.01 1C. NS ¼ Nansen Basin; GR ¼ Gakkel Ridge; AB ¼ Amundsen Basin; LR ¼ Lomonosov Ridge. (b) Same as panel a but now showing profiles from Amundsen Basin only (stations 52–57). Successive profiles are offset by 0.005 1C. ARTICLE IN PRESS 1260 G. Bjo¨rk, P. Winsor / Deep-Sea Research I 53 (2006) 1253–1271 bottom layer. The layer is present in the deep parts and 31 (555 m) further west. The temperature is of the Nansen and Amundsen Basins, extending generally higher during 2001, but the layer is now over the width of the Amundsen Basin, from the thinner at the eastern station 36 (406 m) compared area close to the Siberian shelf break, as observed to station 54 (460 m) and station 60 (525 m). during ARKXII, to the central part of the basin and Fig. 6 shows depth profiles of some selected closer to the Canadian Archipelago. The layer stations, located in close proximity of each other thickness in the Amundsen Basin is between 350 (Fig. 1), from the Amundsen and Nansen Basins, and 700 m, where the 700-m layers are found in the which have been chosen to show temporal develop- ARKXII data (see Table 1). The layer in the ment of the layers. Profiles of potential temperature Nansen Basin is not so widespread and is less and salinity from 1991 and 1996 in the Amundsen distinct than in the Amundsen Basin. Basin (Fig. 6a–c) look relatively similar, with Comparing profiles at different locations along slightly higher temperatures and 50 m thicker the east-west axis of the Amundsen Basin during layer in 1996. In 2001, however, the profile is 0.01– 1991 and 2001 (Fig. 5) shows that the layer is 0.1 1C warmer from the bottom up to about 1400 m thicker in 1991 at the eastern section represented by depth and the thickness of the homogeneous bottom station 17 (625 m) compared to stations 30 (470 m) layer is almost halved, about 400 m compared to

Θ (°C) -0.96 -0.95 -0.94 -0.93 3000

3100

3200

3300 1991 2001 3400

3500

3600

3700

3800 Pressure (dbar) 3900

4000

4100

4200 1991 #17 1991 #30 1991 #31 4300 2001 #36 2001 #54 4400 2001 #60

Fig. 5. Vertical profiles of potential temperature, Y, in the Amundsen Basin. Selected stations along the east–west basin axis in 1991 and all deep stations in 2001. Profiles plotted in black are from the eastern part of the Amundsen Basin, and profiles plotted in dark and light gray are from the western part of the basin. See also Fig. 1 for station locations. ARTICLE IN PRESS G. Bjo¨rk, P. Winsor / Deep-Sea Research I 53 (2006) 1253–1271 1261

Amundsen Basin 1000 1000 3000 (a) (b) (c)

1500 1500 3250

2000 2000 3500 2500 2500 3750 3000 3000 Pressure (dbar) 4000 3500 3500 1991 #17 1996 #54 4000 4000 2001 #36 4250 Makarov 4500 4500 4500 -0.8 -0.6 -0.434.91 34.93 34.95 -0.96 -0.95 -0.94 -0.93

Nansen Basin 1000 1000 3000 (d) (e) (f)

1500 1500 3250

2000 2000 3500 2500 2500 3750 3000 3000 Pressure (dbar) 4000 3500 3500 1991 #05 1994 #37 4000 4000 4250 2001 #26

4500 4500 4500 -0.8 -0.6 -0.434.91 34.93 34.95 -0.96 -0.95 -0.94 -0.93 Θ (°C) Salinity Θ (°C)

Fig. 6. Temporal evolution of temperature, Y, and salinity, S, proﬁles in the Amundsen Basin from 1991, 1996, and 2001 (panels a–c), and Nansen Basin from 1991, 1994, and 2001 (panels d–f). Left and right panels show Y and middle panels S. Also shown in panels a and b is a proﬁle from the Makarov Basin (AO-01 number 43; blue dotted line). Station locations are shown in Fig. 1. The arrow in the (a) panel marks the center of the Makarov Basin water signal.

700 m in 1991 and 1996. There is also an indication sen Basin is at least 0.01–0.015 1C for all three that the bottom water has somewhat higher density stations over the depth range 1500–4500 m. The in 2001 (Fig. 7), but one should be careful when temperature change in the Amundsen Basin indi- interpreting such small changes since the accuracy cates that the deep water may have been renewed, or of the salinity measurements is probably not better at least partly renewed, between 1996 and 2001, an than 70.002. It should be noted that station 36 is issue we will return to later. somewhat anomalous compared to the other two The situation is rather different in the Nansen stations, 54 and 60, during 2001 (see also Fig. 5) Basin (Fig. 6d–f). Here the mixed layer has become with somewhat higher temperature between 1500 gradually thicker when stations from 1991, 1994 and 3000 m and with a more pronounced salinity and 2001 are compared, while the temperature is maximum. The temperature increase in the Amund- slightly lower in 1994 compared to 1991 and then ARTICLE IN PRESS 1262 G. Bjo¨rk, P. Winsor / Deep-Sea Research I 53 (2006) 1253–1271

Amundsen Basin Nansen Basin -0.92 -0.92

-0.93 -0.93

-0.94 -0.94

28.1

28.09 28.1

28.11

C) 28.09 28.11 C) 28.105

28.105 28.095

° 28.095 ° ( ( Θ Θ -0.95 -0.95

-0.96 -0.96

(a) (b) -0.97 -0.97 34.92 34.93 34.94 34.95 34.92 34.93 34.94 34.95 S S

Fig. 7. Deep Y-S profiles (42500 dbar) in the Amundsen Basin from 1991, 1996, and 2001 (panel a), and Nansen Basin from 1991, 1994, and 2001 (panel b). The stations are the same as in Fig. 6. Filled markers show 3000 and 3500 m depth. somewhat higher again in 2001. The temperature in 2001. In connection with a larger water exchange it the Nansen Basin is constant within 0.005 1C over is likely that the bottom mixed layer disappears and the depth interval 2000–4000 m at these three that the bottom water is replaced with slightly stations. The difference between the Nansen and stratified water all the way to the bottom. Such a re- Amundsen Basins with respect to increased tem- setting of the mixed layer can be used to investigate perature is perhaps most clearly seen by comparing the time evolution of a new bottom mixed layer. The Fig. 6a and d. Note that all three stations in the next section describes a simple one-dimensional Nansen Basin (Fig. 6d) merge into one line in the model used to test if it is possible to form a profile depth interval 2000–4000 m, which is not the case like the one observed in 2001 using the observed for the Amundsen Basin (Fig. 6a). The temperature heat flux, starting with a hypothetical, stratified changes in the Nansen Basin may still be significant initial profile. as an indicator of water exchanges, but since the signals here are rather small we focus on the larger 3.1. Time development signal in the Amundsen Basin. Here we investigate the evolution of a convecting 3. Dynamics of the bottom layer bottom layer using a one-dimensional diffusion/ convection model that calculates the time develop- The observations from 1991 and 1996 show that ment of temperature and salinity for a given the well-mixed bottom layer is rather uniform over (constant) geothermal heat flux. The model used each of the deep basins and with no distinct time here is the simplest possible. It should be mentioned development. This indicates that the layer was that much more sophisticated boundary layer nearly in a steady state. It should be kept in mind models exist, for example those of by Large et al. that the sampling interval is rather long (5 years) (1994) and Mironov et al. (2002). However, in the meaning that there exists a possibility for variations present application, with little data available and between observations. A striking feature is that the unknown initial conditions, we do not see the need profiles in the Amundsen Basin are very different in to implement a very detailed model. The model has 2001 (compared to 1991 and 1996), with signifi- a vertical resolution of 1 m, and the computations cantly higher temperatures over a depth range start by heating the lowest grid cell according to the (1700 m) much larger than the well mixed bottom applied geothermal heating, layer, indicating that a relatively large water dT F exchange occurred in the Amundsen Basin with ¼ , (1) supply of warmer water sometime between 1996 and dt r0Cp dz ARTICLE IN PRESS G. Bjo¨rk, P. Winsor / Deep-Sea Research I 53 (2006) 1253–1271 1263 where T is the temperature in the lowest grid level, the top of the mixed layer like the one seen for the F the geothermal heat flux, ro a reference density, salinity. This is the case for all stations with a well- Cp the heat capacity, dt the time step (1 day), and mixed bottom layer. The lack of a density step at the dz the vertical size of a grid cell (1 m). The top of the mixed layer implies that no effect of model column is then mixed upward cell by cell mechanical mixing due to convective motions until a stable profile is obtained (convective adjust- (penetrative convection) can be seen (at least to ment), which also gives an initial mixed layer within the instrumental noise). The model computa- thickness for each time step. Additional mixing tions are therefore made without penetrative con- due to penetrative convection is also implemented in vection. the model, but, as will be shown later, this process Consecutive profiles in Fig. 8 show the model seems to be of minor importance. Diffusive mixing computed time development of the bottom layer is given by during 9 years for two values of the turbulent diffusivity. The actual value of the turbulent dx D i ¼ ðÞ þ diffusion in the Amundsen Basin is not known, 2 xiþ1 2xi xi1 , (2) dt dz and double diffusive processes could be important where x is the variable (temperature or salinity), i such that heat and salt diffuse at different rates. the grid cell index, and D the diffusion coefficient. Typical turbulent diffusivities observed in other The procedure includes splitting of the time step in deep basins such as the Brazil Basin are order to ensure numerical stability. The mixing time 0.11 104 m2 s1 (Polzin et al., 1997). Keeping step is an integer fraction of the overall time step with our simplified approach we use the same and obeys the condition dtpdz2/2D. A notable diffusivity for salt and heat and test for two characteristic of the diffusion process is that it tends different values of the diffusion coefficient (0.2 to re-stratify the upper part of the mixed layer and and 1.0 104 m2 s1). As seen from the tempera- acts to smooth out the edge-like structure at the top ture plot, heating and convective adjustment gen- of the mixed layer, where the properties change erate a well-mixed bottom layer that becomes from being vertically homogeneous to stratified with thicker and warmer with time and the diffusion depth. This will affect a number of depth cells (equal gives a smooth profile above the mixed layer. The to the number of split time steps) in the upper part bottom layer becomes warmer than the observa- of the mixed layer. Once the diffusion step is tions after about 5 years. The salinity in the mixed finished, the procedure starts over again with layer decreases with time because of mixing with less heating and convection (defining a new mixed saline water above the mixed layer, and a step-like layer) and so on. In order to include compressibi- structure develops. The convective adjustment en- lity effects, which become important at high sures that the density profile is stable or at least pressures, the potential temperature and potential marginally stable. Comparing the cases with low density are referred to 4000 dbar in the model and high diffusivity it is seen that increased computations (denoted y4 and s4, respectively). The diffusion decreases the thickness of the mixed layer imposed geothermal heat flux is constant at since diffusion tends to stabilize the upper part of 80 mW m2. the mixed layer. The case with a low diffusivity We use station 36 from AO-01 to exemplify how seems to give the best fit with the observed profile the time development may occur. The numerical and this occurs after 4–5 years from the start of the calculations start from hypothetical parabolic initial simulation. The maximum diffusive heat flux occurs temperature and salinity profiles without any just above the mixed layer and is about 25 mW m2 bottom mixed layer. Fig. 8 shows the observed (after 5 years) for the low value of diffusivity. This potential temperature, salinity and potential density means that about 30% of the imposed geothermal at station 36 (grey lines) as well as initial profiles heat flow leaves the mixed layer. and model results. The initial profile used (0-year The time scale of 4–5 years for development of a line in Fig. 8a) is the same as the observations at well-mixed layer similar to the observations coin- station 36 except for the lowest part, which is cides with the time between the 1996 and 2001 replaced by a parabolic profile. The vertical gradient expeditions. The bottom waters of the Amundsen of salinity and temperature is similar to the Basin must then have been renewed within one or observed profile just above the mixed layer and two years after the 1996 expedition, according to zero at the bottom. Note that s4 shows no step at this model calculation. ARTICLE IN PRESS 1264 G. Bjo¨rk, P. Winsor / Deep-Sea Research I 53 (2006) 1253–1271

3500 3500 3500 0 yrs 3 6 9 2001 #36

4000 4000 4000 Pressure (dbar)

(a) (b) (c) 4500 4500 4500 -0.694 -0.69 -0.686 -0.682 34.943 34.945 46.417 46.42 46.423 Θ ° σ 3 4 ( C) S 4 (kg/m )

3500 3500 3500 0 yrs 3 6 9 2001 #36

4000 4000 4000 Pressure (dbar)

(d) (e) (f) 4500 4500 4500 -0.694 -0.69 -0.686 -0.682 34.943 34.945 46.417 46.42 46.423 Θ ° 4 ( C) S σ 3 4 (kg/m )

Fig. 8. Modeled time development of the bottom layer during 9 years together with observations at station 36 from the AO-01 expedition

(2001). Potential temperature, Y4, and density, s4, are referred to 4000 dbar pressure. The upper three panels (a–c) are for a low turbulent diffusion coefﬁcient D ¼ 0.2 104 m2 s1, and the lower panels (d–f) for high diffusion D ¼ 1.0 104 m2 s1. The 0-year line shows the initial proﬁle used.

3.2. Bottom layer in steady state and temperature surplus and no clear time dependence. One possible explanation is that the layer was Although it is possible to explain the observations in quasi steady state between 1991 and 1996, such in 2001 in terms of a developing well-mixed bottom that the supply of geothermal heat was balanced by layer, this cannot explain the observations from upward diffusion of heat above the layer and heat 1991 and 1996, which show similar layer thickness exchange with somewhat colder boundaries. A ARTICLE IN PRESS G. Bjo¨rk, P. Winsor / Deep-Sea Research I 53 (2006) 1253–1271 1265 likely mechanism for exchange with the boundaries where A is the area of the Amundsen Basin deeper is by thermohaline intrusions (see schematic, Fig. 9). than 4000 m (A ¼ 2 1011 m2). In order to find a Intrusive structures are common in the Arctic, in realistic temperature difference, DT, between the areas where water masses with slightly different boundary and interior we use observations from temperature and salinity but equal density meet (e.g. 1996 to 1991 which show a typical difference of Perkin and Lewis, 1984; Rudels et al., 1999). They 0.01 1C. The boundary is actually warmer than the are, for example, very distinct where the Barents Sea interior in the 1996 observations, but we assume inflow merges with the warmer and saltier inflow of that such a difference is realistic during periods with Atlantic water from Fram Strait, near St Anna a colder boundary. Using DT ¼ 0.01 1C gives a Trough (Schauer et al., 2002). Intrusive features can volume transport of 0.4 Sv for the intrusive motions. also be seen near the boundaries in the 1991 and Distributing this transport over an area with 500 m 1996 sections, for example station 58 ARKXII and height and with the same length as the Amundsen station 18 Oden’91, (Fig. 10) which show well- Basin (1100 km) gives a typical speed of 0.7 mm s1. defined deep-ocean (43200 m) intrusions. The This is a realistic velocity scale for intrusive motions typical signature of intrusive features is that both which typically have velocities about 1 mm s1 or temperature and salinity form distinct maximums or less. minimums at the same depth. The partial density Also indicated in Fig. 9 are dense plumes from differences due to temperature and salinity are the shelf as a likely source of cold water in order compensated such that the profile is stable. Both to keep the boundary at a lower temperature than these stations are situated between a boundary the interior. A problem with sinking shelf plumes in current core and the interior, and have intrusions most Arctic areas in this respect is that they tend that are up to 50 m thick. to entrain warm water during the decent through Using 80 mW m2 for the geothermal heat flux the Atlantic layer such that even if the initial together with estimates of the area of the Amundsen temperature is close to the freezing point, the Basin and the thickness of the well mixed layer, it is final temperature becomes higher than the bottom possible to estimate the water exchange Q needed water temperature. There is however one possible for a steady state as source of cold shelf water at deep levels in the St. Anna Trough as suggested by Rudels et al. AF Q ¼ , (3) (2000). Here, an inflow of cold water has been r0CpDT observed which occasionally might be dense enough to penetrate down to the deep and bottom waters and without entraining much heat. The reason why the heat entrainment should be small for Cold dense current this particular shelf flow is that it enters the Eurasian Basin at a relatively deep level such that Stronger Weaker the entrainment of ambient water will start mixing mixing below 1000 m, which is deeper than the core of Atlantic water. The coldest part of the St. Anna Trough section during 1996 had a temperature of 1.25 1C(Schauer et al., 2002). Using the temperature difference between the St. Anna Colder Warmer 1 Fresher Saltier inflow and Amundsen Basin deep water (0.95 C) gives a volume transport of 1.3 104 m3 s1 in order to keep the boundary water mass in steady state. This is less than 1% of the total transport of Geothermal heatflow cold water in the channel, which was estimated to Fig. 9. Schematic of a possible exchange mechanism for keeping about be 2 Sv during the 1996 expedition (Schauer the warm mixed bottom layer in a steady state. Heat is exchanged et al., 2002). It seems therefore that the inflow between the colder boundary and warm interior through intrusive of cold water from St. Anna Trough can provide motions in both directions. The boundary is kept at a low temperature by cold dense currents coming from the shelf. a source of cold water strong enough to balance Stronger mixing at the boundary will help to transport heat the geothermal heat supply in the Amundsen upward and balance the geothermal heat flow. Basin. ARTICLE IN PRESS 1266 G. Bjo¨rk, P. Winsor / Deep-Sea Research I 53 (2006) 1253–1271

3000 3000 1991 #18 1996 #58

3500 3500 Pressure (dbar)

4000 4000

(a) (b) 4500 4500 0.955 0.95 0.945 0.94 0.935 34.935 34.94 34.94 Θ (°C) S

Fig. 10. Vertical proﬁles of Y and S for station 18 from the Oden’91 expedition and station 58 from the ARK XII (1996) expedition. The salinity proﬁle for station 18 is offset by 0.002 to the right.

4. Evidence of deep-water renewal in the Amundsen neglecting quality issues, they show that the Basin temperature at 4000 m in the Amundsen Basin has ranged between 0.97 and 0.95 1C from the 1950s Our observations from 2001 show that the deep to 1980s, with decadal means being constant around and bottom water below 1500 m in the Amundsen 0.96 1C over four decades. The observed annual Basin is about 0.01–0.021 warmer (depending on mean temperatures from 1991 to 1996 at 4000 m are depth) compared to hydrographic observations in 0.95 1C (when averaged to two decimal places) the 1990s (see e.g. Fig. 6). Historical deep-water and falls within the observed range during the temperatures are available from the EWG Atlas, 1980s. The historical temperature data leads to two grouped into decadal means from the 1950s to the conclusions: First it seems as if the deep water in the 1980s. Fig. 11 summarizes all available deep-water Amundsen Basin has stayed at a relatively constant temperatures from the Amundsen Basin. The temperature between the 1950s and the 1980s, quality of the historical temperatures is question- suggesting that advective processes were, on decadal able as to the accuracy needed to compare with the averages, balancing the geothermal ﬂux, and sec- high-quality CTD data from the 1990s. However, ond, that the jump in temperature between 1996 and ARTICLE IN PRESS G. Bjo¨rk, P. Winsor / Deep-Sea Research I 53 (2006) 1253–1271 1267

-0.94

-0.945

-0.95 1

8 5

-0.955 C Θ ° -0.96 9 12 4

-0.965 3

-0.97

-0.975 1955 1960 1965 1970 1975 1980 1985 1990 1995 2000 Year

Fig. 11. Summary of available hydrographic observations of deep-water temperatures from the Amundsen Basin. Black dots to the left of the dashed line are historical observations of water temperatures at 4000 m from the EWG Atlas given as decadal means. To the right of the dashed line are the deep-water temperatures from the four expeditions presented in this paper (1991, 1994, 1996, and 2001). Vertical lines show the range in temperature from all observations and the digit at each dot shows the number of observations for each decade (1950s–1980s) or year (1991–2001).

2001 is large and unique compared to historical Fig. 6a). The same signal is absent in the Nansen decadal changes. The water temperatures of 2001 Basin. It is, however, not possible to explain the clearly stand out as the warmest during the last five warming below the maximum by inflow from the decades. Makarov Basin since the density is too low to reach Where does this warmer water come from? One the deep Amundsen Basin (Fig. 12). This is further possibility is inflow of warm water from adjacent supported by the salinity data which show that this areas. Fig. 12 gives an overview of the TS signal does not reach below 2400 m. Note also characteristics of the major deep basins in the area. that the density of the Makarov Basin water will be The water in the Makarov Basin water is generally reduced further compared to Amundsen Basin deep warmer and saltier than the Amundsen Basin water water if the density is evaluated at higher pressure at the same depth and is thought to flow mainly since the warmer water in the Makarov Basin has over the Lomonosov Ridge (from the Makarov lower compressibility. Basin to the Amundsen Basin) close to the Another possibility is that warmer deep water Canadian archipelago (Jones et al., 1995). An may have been advected into the Amundsen Basin influence of Makarov Basin water can be seen in from the Norwegian/Greenland Sea. This water hydrographic data from all three years in the must then have passed the Nansen Basin confined to Amundsen Basin, especially in salinity, which shows the shelves of the Barents and Kara Seas without a maximum around 1700–1900 m (see Fig. 6). This affecting the properties in the central Nansen Basin. signal is most pronounced in 2001, with a distinct After the passage of the Nansen Basin it has then to salinity maximum located around 2000 m (arrow in fill up, or at least mix with, the deep parts of the ARTICLE IN PRESS 1268 G. Bjo¨rk, P. Winsor / Deep-Sea Research I 53 (2006) 1253–1271

Fig. 12. Y-S proﬁles from the Nansen Basin (stn. 26), Amundsen Basin (stn. 36) and Makarov Basin (stn. 43) from the Oden 2001 expedition together with a station from the Greenland Sea in 1995 (77.8 1N, 7.5 1E). Filled markers show the 1000, 1500, 2000, 2500, 3000 and 3500 dbar levels.

Amundsen Basin. We have made a survey of all make use of a simplified entrainment model to stations of the World Ocean Data Base in the describe how the dense plume evolves during the northeastern part of the Greenland Sea (WMO descent. Here we use the result of this model and squares 1700 and 1701) that cover stations from the estimate the source of cold shelf water needed to 1980s up to 2001 and have not found any evidence generate a temperature change like the one we of a warmer water mass that could explain the observe in the Amundsen Basin. finding in the Amundsen Basin. The deep water in The basic quantity needed for this computation is the eastern part of Greenland Sea below 1500 m was the entrainment rate or dilution factor. Rudels et al. found to be less saline and colder than the deep (2000) found that a dilution factor of 2 for every waters in the Eurasian Basin for the time period 300 m descent fits the data in the Canada Basin and investigated. A typical station from Greenland Sea use this result in the Eurasian Basin as well. We will (77.8 1N, 7.5 1E) is shown in Fig. 12 for reference. assume that the plume starts at 100 m and then The most likely mechanism for generating a entrains ambient water down to 2500 m, where it warming signal in the Amundsen Basin is supply reach its final characteristics. The total dilution will of dense shelf water that has entrained warm water thus be 16 times over 2400 m. We simplify further during the decent along the slope into the deep by assuming that the plume entrains ambient water water. Dense shelf water is produced by the ice of average properties over the entraining interval. production in coastal polynyas in several areas Suppose that we have a supply of dense shelf water around the Eurasian Basin (Winsor and Bjo¨rk, with initial volume flux Q0 and temperature T0 and 2000). Rudels et al. (2000) point out the area around that dense plume entrains warmer ambient water Severnaya Zemlya as a particularly likely source of with temperature TS (denoted slope water in Rudels warm deep water to the Amundsen Basin. They also et al., 2000). Its final temperature TB (boundary ARTICLE IN PRESS G. Bjo¨rk, P. Winsor / Deep-Sea Research I 53 (2006) 1253–1271 1269 temperature) is then given by the Fram Strait Branch becomes warmer, it is likely that the mixture also becomes warmer, and it is T þ eT T ¼ o S , (4) therefore possible that relatively shallow warming B 1 þ e signals in the Fram Strait branch can be transferred where e is the dilution factor. The volume of water, down to the deep waters by entraining shelf plumes. VB, with temp TB needed to raise the deep-water A number of recent papers indicate an anom- temperature by an amount DTD is alously warm and strengthened inflow of Atlantic Water (e.g. Gunn and Muench, 2001). The apparent DT V ¼ D V , (5) warming seen in the 2001 profiles could originate B T T D B D from a shallow and warmer inflow through Fram where TD and VD are the deep-water temperature Strait. This event must then occurred some years and volume respectively. It is also required here that before 2001, as the advection time from Fram Strait V B V D. By assuming that VB is generated over to this part of the Amundsen Basin is 3–6 years a certain time, T, it is possible to estimate the flux (e.g. Karcher and Oberhuber, 2002). Saloranta and as: QB ¼ VB/T and the initial shelf flow from Haugan (2001) examined hydrographic data from Q0 ¼ QB/e. For the slope water temperature we west of Svalbard, centered in the West Spitsbergen use TS ¼0.7 1C based on an average of the Current. They found that a dominant warming temperature profile at the slope during 1995 event in the upper 300 m occurred at the beginning presented by Rudels et al. (2000). The initial plume of the 1990s, culminating in 1992, with peak temperature should be close to the freezing point at temperatures over 1.0 1C warmer than the 25-year a salinity around 35, and we use accordingly mean. Furthermore, they found another warming T0 ¼1.8 1C. The observed change in deep-water event starting from 1995 and onwards. The timing temperature, DTD, is about 0.01 1C. Using TD ¼ of these events fits well with our observations. 0.94 1C for the deep-water temperature, VD ¼ 2 1014 m3 for the deep-water volume and a timescale 5. Discussion of one year for the exchange (T ¼ 1 year) gives 5 3 1 4 3 1 QB ¼ 3.1 10 m s and Q0 ¼ 2.2 10 m s . The observed 0.01 1C warming of the bottom In order to put the estimated Q0 into perspective waters in the Amundsen Basin between 1996 and it can be compared with an estimate of total 2001 is based on observations using one single polynya produced dense water (with salinity434.8) thermistor at three deep stations in the Amundsen in the Barents and Kara Seas of 3.7 104 m3 s1 Basin. Small changes of deep-water properties may (Winsor and Bjo¨rk, 2000). The estimated Q0 seems also originate from instrumental errors, sensor then somewhat high if it should represent the resolutions and intercalibrations. However, the polynyas around Severnaya Zemlya since it is hard thermistor used in 2001 (Sea-Bird SBE3+) was to believe that this relatively small area is able to new from the factory at the beginning of the produce as much as 2/3 of the total in the entire expedition, and the post-calibration showed a drift Barents and Kara Seas. The result is, however, of only 0.00051 1C, about 20 times smaller than the rather sensitive to the slope-water temperature. observed signal. An extra check of the validity of Using a somewhat higher temperature, TB ¼ the temperature signal can be obtained from the 0.5 1C, gives a more realistic value, Q0 ¼ 1.2 temperature dependence of conductivity, which also 104 m3 s1, which is only about 30% of the total affects the salinity. If the actual bottom-water from Barents and Kara Seas. temperature were 0.01 1C lower than the reading It thus seems possible to generate a 0.01 1C from the thermistor, it would give an unrealistically increase of the Amundsen Basin deep-water tem- high bottom water salinity of 34.956 instead of perature in one year by a supply of shelf-derived 34.944 for the same conductivity (using data from water that has mixed with a warm water column at 4200 dbar at station 32 as an example). This the slope provided that the slope water is somewhat difference of 0.012 is much larger than the accuracy warmer than in 1995. The slope water east of in the salinity data, which is better than 0.002 based Severnaya Zemlya is a mixture of the two main on pre- and post-cruise calibration of the conduc- inflowing branches of Atlantic water: the Barents tivity cell and bottle salinities run on a laboratory Sea Branch (entering at the St. Anna Through) and salinometer. Similar uncertainties in temperature the Fram Strait Branch (eg. Rudels et al., 2000). If are found for the 1991, 1994, and 1996 cruises, ARTICLE IN PRESS 1270 G. Bjo¨rk, P. Winsor / Deep-Sea Research I 53 (2006) 1253–1271 typically being 0.001–0.002 1C or less. Thus, we 6. Concluding remarks believe that the observed temperature increase in the deep water of the Amundsen Basin is real and not The observed well-mixed bottom layer over large an instrumental artifact. parts of the Eurasian Basin implies that, away from It seems clear from the present study that the deep boundaries, much of the deep-water mixing in the waters of the Eurasian Basin are dynamically active abyssal Eurasian Basin is driven by geothermal heat in the sense of having a rather quick response to flux supplied from the Earth’s interior, i.e. a very near-surface processes, although the changes in different mechanism than what is thought to control the deep water are small in absolute terms. An deep-water mixing in the other oceans of the world. active behavior of the deep waters is also indicated The geothermal heat flux decreases the density of from transient tracer studies showing significant the bottom water both by mixing and by direct amounts of bomb tritium (Schlosser) and CFCs heating of the water column. The continuous (Jones et al., 1995) in the deep waters. The basic density decrease by heating is probably important communication link between the surface and abyss for the rate of deep-water renewal by increasing the is maintained by dense shelf plumes bringing sur- likelihood that water masses descending down the face and intermediate water properties down to the slopes will have density high enough to penetrate deep water. into the deep water. The extent to which the deep- It is likely that the geothermal heat flux is water renewal is a continuous or intermittent important for the replenishment of the deepest process seems to be an open question at this stage. layers since, it acts to reduce the density and will Observations prior to 1990 indicate a more con- precondition the water column to supply of new tinuous renewal with relatively constant deep-water (and denser) bottom water. A recent paper by temperatures, while later observations show a Timmermans et al. (2003) confirms that vertically sudden temperature jump indicating one or a few homogenous layers, formed by geothermal heat larger, but relatively short-lived, events. The deep flow, also exist in the Canada Basin. waters seem to have a relatively rapid communica- The average insolation age based on 14C mea- tion with the surface, and dense bottom currents surements is 250 years in the Eurasian Basin deep generated at the shelves might play an important water with a slightly shorter time scale in the role for the exchange. Amundsen Basin compared to the Nansen Basin. A relevant question is then how such a relatively long time scale fits with the short time frame of o5 years Acknowledgements for the observed temperature increase in the Amundsen Basin between 1996 and 2001. The We thank N. Kaul for generously sharing answer is that the temperature increase is probably unpublished heat flow data. We also thank L. caused by supply of a small volume of surface Arneborg and three anonymous reviewers for many water, which has become almost completely mixed constructive comments on our paper. Funding for with the entire deep-water volume. This means that G. Bjo¨rk came through the Swedish Research this type of renewal will give only small changes in Council (contract no: 621-2001-2578). Financial most constituents such as D14C that are therefore support was provided to P. Winsor from NSF hard to detect. Using the same observed D14C OPP-0352628. We are grateful for this support. profiles as in Schlosser et al. (1997) with a surface value of about 80%, bottom value of 80% and an average between 100 and 2500 m of about 0%,itis References possible to estimate the expected change of D14Cin the deep water by using the simple mixing model 4 3 1 Aagaard, K., Coachman, L.K., Carmack, E., 1981. On the (Eq. (4) and (5)). Using Q0 ¼ 1.2 10 m s for halocline of the Arctic Ocean. Deep-Sea Research 28A, the initial shelf flow, it turns out that the same water 529–545. renewal that increased the deep-water temperature Aagaard, K., Swift, J.H., Carmack, E., 1985. Thermohaline by 0.01 1C will increase the D14Cby2.5%. Such circulation in the Arctic Mediterranean Seas. Journal of Geophysical Research 90, 4833–4846. a change might be possible to detect since the 14 Adcroft, A., Scott, J.R., Marotzke, J., 2001. Impact of precision of the D C measurements is typically geothermal heating on the global ocean circulation. Geo- 72–5% (Schlosser et al., 1997). physical Research Letters 28, 1735–1738. ARTICLE IN PRESS G. Bjo¨rk, P. Winsor / Deep-Sea Research I 53 (2006) 1253–1271 1271

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