Oceanic thermal structure in the western Canadian Arctic

Conhnental Shelf Research, Vol 3, No_ 3, pp 233 to 258, 1984 Printed in Great Britain

0278~1343/84 $3 00 + 0 00 {~ 1984 Pergamon Press Ltd

Oceanic thermal structure in the western C a n a d i a n Arctic HUMFREY MELLING,* R. A. LAKE,* D. R. TOPHAM* and D. B.

FISSEL~"

(Received 7 October 1983; m rewsed form 8 February 1984; accepted 10 February 1984)

Abltraet--Recent hydrographic data (1981-1982) from the western Canadian Arctic Archipelago and adjacent areas of the Arctic Ocean are interpreted from the viewpoint of thermal energy transfer. Within the Archipelago, a warmer halochne than m the Arctic Ocean and a cooler Atlantic layer are identified. The warmer haloclme is a consequence of the continued diffusion of heat from underlying Atlantic water without a significant downward penetration from the surface of cold (<--I 5oc) seawater with salimty increased consequent to ice growth The cooler Atlantic layer is primarily attributable to an enhanced coohng of these waters m a narrow band over the continental slope and shelf of the southern Beaufort Sea prior to their inflow into the Archipelago Rates of transport and vertical dMuslon in this region are estimated The significance of these find.rigs in regional and Arctic oceanography is discussed

INTRODUCTION

THE C a n a d i a n Arctic Archipelago lies on the extensive polar continental shelf of North A m e r i c a (Fig. l). Roughly triangular, it is bounded on the west by the C a n a d a Basin of the Arctic Ocean and on the east by Baffin Bay. F r o m the oceanographic viewpoint, the Archipelago has western and eastern parts which are separated by the shallow sills (~<100-m depth) found principally in those channels near 90 to 1 0 0 ° W (Fig. 1). These sills prevent flow of all but near-surface waters from the western part into Baffin Bay. A second series of deeper sills (350 to 450-m depth) near the western edge o f the continental shelf limits the inflow o f waters to the Archipelago from the Arctic Ocean. M a x i m u m depths between the western and central sills are typically 550 to 650 m, although depths exceeding 900 m occur in N a n s e n Sound. A sill of 250-m depth in Nares Strait m a r k s the deepest connection to Baffin Bay, through the Archipelago. Observations relating to the physical o c e a n o g r a p h y of the western C a n a d i a n Arctic Archipelago and adjacent areas o f the Arctic Ocean have been few in number (for documentation, see BIRCH et al., 1983; FlSSEL et al., 1983). Such that exist have permitted the development of a general descriptive o c e a n o g r a p h y o f the western Archipelago (BAILEY, 1957; COLLIN, 1963; COLLIN and DUNBAR, 1964; FORD and HATTERSLEY-SMITH, 1965; BARBER and HUYER, 1971). O c e a n o g r a p h i c understanding of the Arctic Ocean is better developed, and is summarized most recently by TRESHNIKOV and BARANOV (1972) and COACHMAN and AAGAARD (1974). In simplified terms, ice and cold, low salinity surface waters (,~2 x l06 m ~ s -] in total) originating as ice melt, river runoff, and inflow from the Bering Sea are swept by prevathng winds across the central Arctic Ocean into the East * Frozen Sea Research Group, Inshtute of Ocean Sciences, P O Box 6000, Sidney, B.C., Canada V8L 4B2. 1"Arctic Sciences Ltd., 1986 Mills Road, R R No 2, Sidney, B C., Canada VBL 3SI. 233

234

H. MELLINGe/al.

90"N

80" N

75" N

70Jl~

110"W

Fig. I. Bathymetnc chart of the western Canadian Arctic (based on GEBCO map 5-17). Depths (m) of the major sills are indicated (O). Depth contours are as follows: 200 m ( ), 400 m (--), 1000 m ( ), 2000 m (---), and 3000 m (-..). Water bodies are identifiedby letters. Baffin Bay (A), Lancaster Sound (B), Barrow Strait (C), M'Clintock Channel (D), Viscount MelvilleSound (E), Amundsen Gulf (F), M'Clure Strmt (G), Prince Gustaf Adolf Sea (H), Peary Channel (1), Nansen Sound (J), Lincoln Sea (K), Nares Strait (L), Penny Strait (M), and Byam Martin Channel (N). The region between H, I and N, M is frequently referred to as the Sverdrup Basra. G, E, C, and B constitute Parry Channel. The Queen Ehzabeth Islands are those to the north of Parry Channel.

Greenland Current. A n outflow of similar magnitude is thought to pass through the channels of the C a n a d i a n Arcttc Archipelago in response to forcing which is not understood. To compensate for these losses, a flow of warmer, more saline water from the Norwegian Sea (termed Atlantic water) enters the Arctic Basin as the West Spitsbergen Current, descends to depths between 200 and 1000 m north of Spitsbergen and progresses eastward as a subsurface flow along the Eurasian continental slope. In the C a n a d a Basin the surface flow is also wind-driven, forming the anticyclomc Beaufort gyre, and is directed southwestwards along the western b o u n d a r y of the Arctic Archtpelago. The movement of the submerged Atlantic layer in this area, referred from the horizontal distribution of m a x i m u m temperature m sporadic and imprecise hydrographic studies, is weak and all-defined. Although Atlantic water temperature generally decreases in a cyclonic sense a r o u n d the pole, and was first thought to continue in the same sense in the southern C a n a d a

Oceamc thermal structure in the western Canadian ArcUc

235

Basin (TIMOFEYEV,1957), this simple pattern was refined by COACHMAN and BARNES (1963) to include an anticyclonic gyre south of about 80°N in the Canada Basin. Consideration of a more detailed data set by NEWTON and COACHMAN (1974) suggested additionally the existence of a narrow eastward flow of Atlantic water along the continental slope of the Chukchi Sea to the south of the anticyclonic gyre. Hydrographic profiles in the Canada Basin adjacent to the Canadian Archipelago reveal a surface mixed layer ,,-l0 to 50 m in depth and an underlying uniformly cold ( - 1 . 5 to - 1 . 8 °C) layer of increasing salinity extending to ,,-200 m, below which a sharp thermocline marks the transition to Atlantic waters which have a temperature here of 0.4 to 0.5°C at 450 to 500 m. Atlantic waters occupy most of the remainder of the upper 1000 m. The intermediate cold layer constitutes the 'cold pycnocllne' of the Arctic Ocean, a hydrographic feature that has long intrigued oceanographers. It is now apparent that the low temperature of the Arctic pycnocline is attributable to lateral intrusions of suitably saline water at freezing temperature (GARRISONand BECKER, 1976; AAGAARDe/al., 198 l; MELLING and LEWIS,1982). Such water is generated at the periphery of the polar pack during the formation of the seasonal sea-ice cover in winter. Because ice forms on open water at a greater rate than under pre-existing ice, any persistent divergence of a sea-ice cover (with concomitant formation of leads and polynyas) greatly increases the wintertime formation of cold saline water. In the southeastern Beaufort Sea, the Chukchi Sea, and the seas of the Siberian Shelf, the prevailing winds force a net divergence of ice cover (HIBLER,1979), thus favouring these areas as source regions for subsurface water. The broad continental shelves which underlie these shallow seas also play an important role. Since their depth is comparable to the depth of the surface mixing layer (,,~50 m), and since the Coriohs effect strongly retards flow into deeper areas, waters are retained near the freezing interface for long periods and the large salinity increase (>~2) required for penetration into the pycnocline is thus possible. M ELLINGand LEWIS(1982) estimate such retention times to be the order of 2 to 6 months. If the water depth were appreciably greater than that of surface mixing, surface water with salinity only slightly enhanced by freezing would descend beneath the realm of surface freezing before an appreciable salinity increase (>2) had occurred (MCPHEE, 1980). The temperature-salimty relationship within the Arctic pycnochne is the consequence of a balance between this injection of wintertime shelf water, and the effects of an upward diffusion of heat and salt from the underlying Atlantic water, of an export of Arctic water via the East Greenland Current and through the Canadian Arctic Archipelago (AAGAARDand GREISMAN, 1975), and of an upward entrainment of salt into the river-freshened surface layer. Historical observations within the western Archipelago reveal a modified version of this Arctic Ocean hydrography. BAILEY (1957) notes that in 1954 water in Viscount Melville Sound of salinity 33 to 34.5 (the pycnocline) was warmer than that in the Beaufort Sea by 0.2°C. HUYER and BARBER (1970) show by mapping temperature at a salinity of 33.8 for 1961 and 1962 that pycnocline temperatures were higher than Arctic Ocean values throughout the Archipelago in these years also. M UENCH(1971) has hypothesized that this warmer halocline found in the western and northern Archipelago results from a mixing, by an unspecified process, of waters from the Ameraslan and Eurasian basins of the Arctic Ocean. In this paper modern hydrographic data obtained by profihng conductivity-temperaturedepth (CTD) probes are utilized to investigate the thermal modification of waters originating in the Arctic Ocean as they move over the continental shelf and slope on its North American

236

H . MELLINGet aL

periphery and into the main channels of the western Archipelago. Several simple models are used to illustrate advective--diffusive heat transfers. The implications of these processes for Arctic oceanography are discussed. FIELD

PROGRAMS

Hydrographic measurements by CTD probe are the basis of this study. A total of 187 stations were occupied in the March/April period, 54 in the Beaufort Sea in 1981 and 133 in the Arctic Archipelago in 1982. Some use is made of 12 CTD stations in the Beaufort Sea occupied in late November 1979 (MELLING, 1983). Observations were made by deploying the CTD probe (Guildline Model 8706) through sea ice from an aircraft-transportable winch. Both fixed-wing aircraft (DHC-6 Twin Otter) and helicopters (Bell 206-L 1; Sikorsky $6 IN) were used for transportation and sea-ice landings. Aircraft were positioned using VLF/ Omega radio navigation (e.g., GNS 500A) to a typical accuracy of + 1 km (worst ,,,2 km). The CTD probes were calibrated in the field to establish observational accuracies of +0.005°C in temperature, +0.010 in salinity, and max (0.3 dbar, 0.2%) in pressure. Measurement precision is 1/3 to 1/5 of these values. Precision thermistors (+0.002°C) referenced to a triple-point bath were used as transfer standards in temperature calibration, while water samples analyzed on a Guiidline Autosal 8400 salinometer (+0.001 in salinity) provided reference data for conductivity ratio. A deadweight tester was used to calibrate at pressures above 100 dbar, while at lower values pressure computed at measured depth was the reference. Following correction for cell constant variation with temperature and pressure, and for the different impulse responses of the temperature and conductivity sensors, salinities (quoted without units) were calculated using the Practical Salinity Scale, 1978 (LEwiS, 1980). ARCTIC

WATERS

OF THE PYCNOCLINEIN

THE WESTERN

ARCHIPELAGO

The temperature-salinity (T-S) characteristics of waters within the western Archipelago are shown in Fig. 2, with curves from adjacent areas of the Canada Basin included for comparison. Station positions are mapped in Fig. 3. Figure 2a is a composite T - S diagram for a line of stations extending from the Beaufort Sea west of Banks Island to M'Clintock Channel. The Beaufort Sea profile, which extends to 1000 dbar over the continental slope, displays the usual temperature maximum in the Atlantic layer at 475 dbar (0.41°C), while the temperature near the mid-point of the halocline (at a salinity of 33.5 and pressure of ! 45 dbar) ~s - 1.45 °C. Within Parry Channel, the waters with salinity above 34.76 are uniform in their T - S relationship but significantly cooler than those of the same salinity in the Beaufort Sea. Above the sill depth waters within the halocline warm progressively towards the east. The temperature increase is greatest at a salinity of 33.5 amounting to 0.25°C for waters which have reached the northern end of M'Clintock Channel (Sta. M4). There is no indication of a penetration of surface waters of freezing temperature deep into the halocline such as observed over the continental slope of the Beaufort Sea (M ELLIN6and LEWIS, 1982). Figure 2b is an analogous T - S diagram for a more limited set of stations among the Queen Elizabeth Islands (see Fig. 3 for positions). Unfortunately, the only observations over the continental slope of the Arctic Ocean in this area were made in 1960 (COLLIN, 1961), and the only observations in the Prince Gustaf Adolf Sea in 1979. Nonetheless a systematic warming towards the southeast of waters within the halochne is suggested and the waters of Atlantic origin (T > 0°C) are cooler than in the Canada Basin here also.

237

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Fig. 2 Composite T - S diagrams for lines of hydrographic stations extending into the Canadmn Archipelago from the west. Note the eastward warming within the halochne and the cooler deep waters (high sahnity) within the Archipelago. The '+'s represent freezing temperature at atmospheric pressure Stations are charted m Fig. 3 (a) Parry Channel statmns are listed m order of decreasing temperature at a sahnlty of 33 5 (that is, east to west). (b) Prince Gustaf Adolf Sea/Penny Strait: statmns are hsted in order of decreasing temperature at a sahmty of 33.5 (that Is, southeast to northwest). (c) Amundsen Gulf. (d) Lincoln Sea.

Figures 2c and d depict differences in T - S relationship which are observed between waters in Amundsen Gulf (MELLInG, 1983) and in the Lincoln Sea (SEBERT, 1968) when compared w~th waters in adjacent areas of the Arctic Ocean. Again a warmer halocline and cooler Atlantic waters are evident. In Nansen Sound, observations discussed by FORD and HATTERSLEY-SMITH(1965) reveal similar conditions. A deep coohng and a halocline warming

238

H. MELLINGetal. 90" N

80"h

7S'N

70o~

110~W

Fig. 3. Locations of oceanographic stations utilized in the discussions. Dots correspond to observations in Fig. 2 and circles to those in Fig. 6. Arrows connect with c o m p a n s o n stations m the Canada Basin.

thus seem generally to accompany water movement into the Archipelago from the Canada Basin. The data from Amundsen Gulf in November 1979 (Fig. 2c) are apparently representative of halocline temperatures in the July to November period (MELuNG, 1983). However, they may not be typical of conditions in late winter when cold saline waters generated by freezing over the Mackenzie shelf to the west can intrude into and drastically cool the halocline over the entire gulf (FIssEL et al., 1984). Presumably because of the high rate of inflow of these waters a progressive warming of the halocline may not be evident at this season. By examining temperature on an isohaline surface, spatial patterns in T - S relationship may be investigated. Figure 4a depicts the temperature of water with salinity 33.5 (p ,,~ 145 dbar), and is representative of the waters of the halocline. There Is a large and systematic variation of temperature on this surface. In western M'Clure Strait, near the southern shore, the temperature equals that found over the continental slope in the eastern Beaufort Sea. A flow from the eastern Beaufort Sea into southern M'Clure Strait at this level is thus suggested. The pattern of the isotherms in Fig. 4a indicates that this eastward flow occurs as a current close to the southern shore of Parry Channel. As distance from the Arctic Ocean increases, the

Oceame thermal structure m the western Canadian Arctic

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Fig 4 Temperature (°C) on two lsohahne surfaces m the western Canadian Archipelago m March to Aprd 1982. (a) Temperature at a sahmty of 33.5, representatwe of variations m T - S relationship within the halocline. This parameter, which is typically <-1.45°C m the Beaufort Sea, shows a systematte increase w~thm the Arehtpelago. Data from 1979 (PECK, 1980b) have been utdlzed to posit,on the dashed contours m the Prmce Gustaf Adolf Sea (b) Temperature at a sahn,ty of 34 83, representative of variations ,n T - S relationship within the Atlant,c layer Temperature varmtlons at this level (~400 m) are small

w a t e r s in this c u r r e n t w a r m . W t t h i n the Q u e e n E h z a b e t h I s l a n d s , a w a r m i n g w i t h d i s t a n c e f r o m the A r c t i c O c e a n is also a p p a r e n t f r o m Fig. 4 a b u t a d e f i n i t w e flow p a t t e r n c a n n o t be referred. T e m p e r a t u r e a t a g w e n d i s t a n c e f r o m t h e A r c t i c O c e a n ~s 0.1 t o 0 . 2 ° C h i g h e r a m o n g the Q u e e n E l i z a b e t h I s l a n d s n o r t h o f P a r r y C h a n n e l t h a n in P a r r y C h a n n e l ~tself. T h i s d i f f e r e n c e m a y reflect a d , f f e r e n c e in s o u r c e w a t e r p r o p e r t i e s for t h e t w o r e g i o n s o r a d~fference m t h e heating and advection rates.

240

H. MELL]NO et al. S.E F L O W

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Fig. 5. Profilesof barochmc flow relative to the surface near the northern and southern shores of eastern M'Clure Strait m March to April 1982. Since current at 18 m averaged about 1.5 cm s-' over 10 days following hydrographic observation, the geostrophic flow is primarily baroclimc. SOUTH-I (1.3 to 4.2 km from south shore); SOUTH-2 (4.2 to 9.7 km from south shore); NORTH1 (0.5 to 3.7 km from north shore). The hydrographic evidence for a narrow eastward current within the halocline in Parry Channel (Fig. 4a) is supported by dynamic calculations. Significant baroclinic flow occurs only in a zone extending about one Rossby internal deformation radius (,,-10 k m ) f r o m the southern shore (Fig. 5). On the southern side maximum speeds occur in the middle halocline (p ,,, 150 dbar, S -,, 33.6) in a southeastward flow, relative to the surface. On the northern side speeds within the halocline are low. Between late March and mid-June 1982, near-surface (18 m) currents were measured by the authors at 7 locations across M'Clure Strait using recording current meters suspended from the ice. Over the 10 days of measurement following occupation of the hydrographic section, flows averaged 1 to 1.5 cm s -j towards the east. Such relatively weak surface currents indicate that the baroclinic flow component dominates the geostrophic flow. In the deep channels of the Queen Elizabeth Islands, available data from 1979 (PECK, 1980b,c) and 1982 (FISSELet al., 1984) indicate weak (<1 cm s -l) surface currents and weak geostrophic shears. If narrow nearshore currents m the halocline do exist, they were not resolved by the widely spaced oceanographic stations in these years. Across Penny Strait and on the western side of Byam Martin Channel, both measured surface current and geostrophlc shear indicate strong (10 to 20 cm s -~) southeastward flows. However, because sills to the southeast of the channels restrict transport of water to depths <80 to 100 m and typically to salinities <33, these flow speeds are not representative of the transport of water in the midhalocline (salinity ,--33.5) which is the present interest. Since the progressive warming of inflowing waters within the halocline m southern Parry Channel occurs without the lateral intrusions of freezing-temperature water common in the Beaufort Sea, temperature changes may be modelled by the steady advective-diffusive balance: aT a2T a2T u--= K + K'-(1) ax

~

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Oceanic thermal structure m the western Canadmn Arctic

241

where u is the along-channel flow (eastward), ~T/~x is the eastward"temperature gradient, K is the vertical diffusivity, K ' ts the cross-stream diffusivity, and z ~s the verUcal co-ordinate. A similar equation holds for salinity, S. Since the 33.5 isohaline in 1982 was level across M'Clure Strait, the factor ~2S/~y2 in the salinity equation is effectively zero. Moreover, ~2S/~z2 was so small at this level (,-,10 -5 m -2) that even a verttcal diffuslvity as large as 10 -3 m 2 s -t would result in a negligible change in the salmity of a fluid parcel after advection along the full length of the channel. Thus the advective motion ~s essentially isohaline and the above equation for temperature may be evaluated at constant salimty. Moreover, in Parry Channel it appears that all warming of the halocline ~s ultimately due to heat transport from below. Consequently lateral diffusion will be neglected (the along-channel orientation of isotherms m Fig. 5a does mdicate that cross-channel m~xing is relatively weak). Using values from the 1982 study (u ~ 0.06 m s -~, ~T/~x ~ 0.04°C per 100 km, and ~2T/~z2 3 x 10 4 ° C m-2), the vertical diffusivity, K, is calculated to be 8 × 10-5 m 2 s-~ , a value which is 500 ttmes greater than molecular diffusJvity. Since the largest values of verUcal temperature gradient in the lower thermocline are typically 0.015 °C m -1, this thermal conductivity enables a flux of heat from the deeper Atlantic waters of ~ 5 W m -2 (+50%). For comparison, an average upward heat flux from the Atlantic water over the Canada Basra of 7 to 11 W m -2 (5 to 8 kcal cm -2 y-i) was estimated by AAGAARD and GREISMAN (1975) on the basis of a heat-budget calculataon for the Arctic Ocean. Beyond the confines of the eastward stream there was no appreciable increase m the temperature of the halocline towards the east in the spring of 1982, although ~2 T/~z 2 was not appreciably smaller beyond the stream than within it. Reference to the equatton representing the advecUve-diffusive balance shows that the vertical diffuswity must as a consequence be considerably smaller beyond the stream. A relaUonsh~p between the energy of the mean flow and the intensity of vertical mixing is thus implied. There exists a potentml source of energy for m~xmg in the vertical shear of the geostroph~c flow (the vertical gradient of curves in Ftg. 5). However, the large gradient Richardson number in the geostrophic flow (,~100) precludes extraction of mining energy from the shear by dynamical instabihty. ARMI (1979) argues that the apparent vertical diffuslvity m the deep ocean may be a reflection of an eplpycnal transport of m~xed fluid into the ocean interior following mixing of straUfied fired in the boundary layers formed where isopycnals intersect the sea floor. The vertical dlffUSlvlty associated with thin process is h2/x, where h is the vertical displacement of a fluid parcel in a mixing event (h ,,, 1/2 mixed-layer depth) and x is the interval between events. Boundary mixing through this mechamsm over the sloping s~dewall of the channel is a plausible cause of the relatively high diffusivity observed m the undercurrent in Parry Channel. If the undercurrent is meteorologically forced m the Beaufort Sea, x ~ 5 d, and h may be estimated to be 5 m. Dtffusivity is computed to be h:/'c ~ 6 × 10 5 m 2 s-~, a value close to that computed from observations (8 × 10 5 m 2 s ~). The existence of this eastward underfiow in western Parry Channel requires some discussion. Because the ice cover m this region is landfast between November and July, momentum cannot be imparted locally to the ocean by winds in winter. Moreover, since the flow penetrates a basra which allows no exit to the east of these depths (>100 m), it cannot be responding to a favourable pressure gradient, but must m fact work against the pressure gradient which drwes the westward outflow required by continuity. Thus the current must enter the Archipelago with sufficient momentum to allow its passage deep into Parry Channel.

242

H. MELLINGet al.

In 1982, the 33.5 isohaline sloped upward by 10 to 20 m between 120 and 110°W, requiring the inflow ascending on this surface to decelerate by ~ 10 cm s-J between these meridians. Furthermore, reduction in the core speed of the flow arises from the vertical diffusion of momentum within the eastward stream which spreads momentum over a larger volume. Under assumptions discussed earlier, the reduction in core speed by diffusion can be approximated by: i)u u --

~x

~:u ", Km

/)z 2 '

(2)

where Km is the vertical diffusivity for momentum. For a Gaussian profile of current based on Fig. 5 and a'value for K m 10 times larger than thermal dlffusivlty, this equation indicates that at most a hatvmg of inflow speed would occur over ,,-600 km (10K is a reasonable upper bound on K i n ; CRAWFORO, 1982). A comparable deceleration rate is achieved through frictional interaction with the sea floor. Thus the momentum of the inflowing water does appear to be sufficient to carry it the observed 800 km from the shelf break west of M'Clure Strait to 105°W (Fig. 4a) without further forcing, despite the slowing effects of upslope motion, dlffUSlOnal thickening, and boundary fricUon. A T L A N T I C W A T E R S IN W E S T E R N P A R R Y C H A N N E L

Figure 6 provides a detailed look at the T - S relationship of the Atlantic water (T >~0°C) in the western Canadian Arctic. While salinities close to those at the Atlantic temperature maximum in the eastern Canada Basin (34.86) do appear in the western Archipelago, the corresponding temperatures are lower by up to 0.1 to 0.2°C. This anomaly is present in all the western basins (Amundsen Gulf, Parry Channel, Prince Gustaf Adolf Sea, Nansen Sound, and the Lincoln Sea). T - S curves cross near a salinity of 34.7, below which value Archipelago waters are warmer (Fig. 2). To typify spatial variations in Atlantic waters, temperature on the 34.83 isohaline surface has been mapped in Fig. 4b. Temperature variation in Parry Channel, though systematic, is small m comparison with the difference of almost 0.1 °C encountered over the short distance separating the study area from the Canada Basin. At the depth of this isohaline (,,-400 m), warmest waters lie on the north side of Parry Channel. The reduced Atlantic water temperature within the Archipelago may be entirely a consequence of a local removal of heat to warm overlying waters w~thm the halocline as described in Section 3. In western Parry Channel, waters are cooler than those in the Canada Basin at sahnitles exceeding 34.76 (depths below the 330-dbar level). Using the hypsometry of this area, and the 1982 hydrography, the heat deficit below the 330-dbar level, relative to nearby waters in the Canada Basin, is calculated to be 1.5 × 10 j8 J. In 1982 the total heat gain within the inflowing current, integrated between 100 and 280 dbar (equivalently between 33.1 and 34.6 in sahnlty), is 2.1 x 10 a J m -2, and the transport of the current estimated from hydrographic data is about 0.1 x 10 6 m 3 s -j. Thus the heating rate m the inflow is about 1 x l0 j~ J s -j. If this heat were derived from an unreplenished reservoir below the 330-dbar level, the observed heat defictt of Atlantic waters in Parry Channel could accumulate in about 6 months. Supposing that the interval between renewals of deep waters in Parry Channel is comparable to or exceeds this interval, then hydrographic observations made over the last three decades should reveal deep-water temperatures ranging as high as Beaufort Sea values, and

Oceamc thermal structure in the western Canadian Arctic

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t Fig. 6. Comparison of the T-S relaUonshlps characterizing Atlantic waters (T >~0°C) wathm the Archipelago (hght lanes) and m the eastern Canada Basin (heavy lines). CTD staaons wathout symbol adenUfieationare 31 (1981), D2, B1, E5 and 6 (1979) m order of decreasing temperature at a sahmty of 34 8. Note the lower temperature of Archipelago waters at high salimty. Some of the apparent secular varmtion m waters of the Canada Basin may be a reflecaon of observational inaccuracy (+0.015 °C, +0 015%o) m the older hydrographic data

perhaps lower than the values observed in 1982, depending on the time elapsed since the last renewal event. To test this supposition the average temperature and salinity measured between 425 and 475 m in each cruise year [BIRCH et al. (1983) tabulate available data] have been plotted in Fig. 7, with T - S curves for the Beaufort Sea (1981) and Parry Channel (1982) shown for comparison. The statistical uncertainties o f the bottle d a t a ( + 0 . 0 1 5 ° C , +_0.015 in sahnity), due chiefly to measurement inaccuracy, are indicated by the large square in the figure. With this degree o f uncertainty, the bottle d a t a cannot be considered significantly different from the 1982 observations in the same area. Thus the supposition that deep waters are cooled after entering the Archipelago and that they are completely replaced at intervals by Atlantic waters from the C a n a d a Basm is incorrect. The difference between the historical pomts and the curve for Beaufort Sea water on the T - S d i a g r a m in Fig. 7 is significantly greater than observational uncertainty. Over 30 years, water deep in P a r r y Channel has varied very little in temperature and salinity and has always been cooler than water o f the same salinity in the Beaufort Sea. The only reasonable interpretation o f the available observations is that a modified Atlantic water is moving e a s t w a r d from the deep continental shelf into M ' C l u r e Strait more or less continually, and that the temperature o f this water is already within a few hundredths o f a degree of that within P a r r y Channel (0.3 °C). To maintain the steady deep-water characteristics which are observed in western P a r r y Channel (Fig. 7) in the presence of a continual inflow to the b o t t o m of the basin, the pred o m i n a n t d y n a m i c a l balance within the deep waters ( p > 320 dbar) must relate vertical diffu-

244

H. M ELLINGet al.

o L034 .7"7 ,

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Fig. 7. T - S characteristics of the deep waters in western Parry Channel over the last 3 decades. Modern CTD data are shown as curves, where the small box represents observational accuracy. Points for bottle data depict the average value from each crume(O), where the large box represents the extent of the scatter, probably largely expcnmental. A CTD profile from the Beaufort Sea (198 I) is shown for comparison. Stabons axe listed m order of decreasing temperature at a sahmty of 34.82. Numbers represent depths and the bracketed numbers, the year of observation. Year-to-year vanauons in T - S values are seen to be small relauve to changes over the western sdl.

sion to vertical advection. Above the (presumably thin) inflowing layer at the sea floor and below the level of the return flow, the one-dimensional balance is:

aT w ~z = K .

~T ~z 2 ,

(3)

where K , is a thermal diffusivity appropriate to depths in this range. An analogous ¢quatton describes salinity. Integration generates an expression involving fluxes: w+

K

,

~ ln,4~

/)z ]

(Ta-T)=-K,

~T

~)z ;

z~<-320m,

(4)

where T a is the inflow temperature (bottom temperature) and A the basm area at depth. N o t e that the upward increase in the area of the basm, giving a larger diffusion surface, acts like a fictitious advective velocity in the left side o f the equation. Assuming that all the heat accumulating in the halocline (,,,l x 10JJJ s-') is transported diffusively through the upper level (taken here at 320 m), then the right side of the equation yields K , ,,, l x l 0 -4 m 2 s -~ with use o f 1982 data. This value of diffusivity is associated with static stabilities, expressed in terms of Brunt-Viiisiiil~i frequency, which decrease from 3 to I x l 0 -~ s -j between this depth and 450 m. F r o m the left side of the above equation vertical velocity is evaluated as ,-,15 c m d -~. The corresponding vertical transport of 0.1 x l06 m ~ s-' could be provided by inflow at 2 c m s-' withm a 25-m layer across the western sill.

Oceanic thermal structure in the western Canadian Arctic

245

The vertical diffusivity for salt estimated assuming a local advective---diffuswe balance below sill depth and using the computed vertical velocity is about 75% of the value of thermal dlffuswity K,. Within the accuracy of this calculation, the thermal and salt diffusivities in this depth range may thus be considered equal. Observations which might reveal the relatively cool (0.3°C at 34.83) Atlantic water mass hypothesized earlier to occupy the deep (300 to 400 m) shelves which adjoin the western coast of the Archipelago are scarce. In the area northwest of Ellef Ringnes Island in 1960 (CoLHN, 1961), a station over the slope (1239 m sounding) revealed water of temperature 0.41°C at a salinity of 34.83, whereas over the nearby shelf two stations (500-m soundings) showed temperatures of 0.32 to 0.35°C at the same salinity. West of M'Clure Strait in the same year, stations occupied by the USS Burton Island mdicated temperatures on the shelf (390-m soundings) to be cooler than 0.30°C at this salinity, while m the basra to the west a value of 0.37°C was observed (MARINE SURVEYSDIVISION, 1964). In the southern Beaufort Sea, a more complete picture is available from the 1981 study. Although the continental shelf in this area is too shallow (0 to 100 m) to be flooded by waters of sahmty 34.83, the hydrographic measurements made m 1981 clearly show a very narrow band of anomalously cool Atlantic water overlying the steep continental slope of the North American mainland. This band is evident on the map of temperatures on the 34.83 isohahne surface in the Beaufort Sea (Fig. 8) as an abrupt westward displacement of the isotherms over the continental slope. Temperatures in 1979 over a less extensive area of the Beaufort Sea define a similar pattern of isotherms (MELUNO, 1983).

150 °

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Fig. 8 Temperature on the isohahne surface 34.83 in the Beaufort Sea in March to April 1981 The map displays a systematic decrease ifl Atlantic water temperature from west to east which is greater over the continental slope A cyclonic circulation of Atlantic water in this area may thereby be inferred. A stepped profile suggestive of double-diffusive convection was observed at the circled stations Depth contours moving northward from Alaska are 50, 200, 400, 1500, 2500, 3500, and

3800 m

H. MELLING et aL

246

CIRCULATION

AND

MODIFICATION

OF ATLANTIC

WATERS

NEAR

THE

CONTINENTAL

MARGIN

A progressive cooling from west to east is evident on the 34.83 isohalin¢ surface in Fig. 8, both in the deep waters of the Canada Basra and over the continental slope. This trend demonstrates that the movement of Atlantic water in the southern Canada Basin is in fact cyclonic about the pole and in the opposition to the westerly surface flow. Presumably the eastward flow of Atlantic water is continuation of the slope current in the western Beaufort Sea described by NEWTON and COACHMAN(1974), although its width in the east (>250 km) is considerably greater than in the west (,,-100 km). The dynamic topography at the core of the Atlantic layer (450/1000 dbar) deduced from the 1981 observations (Fig. 9) indicates an eastward baroclinic flow of about 0.5 cm s-j. This circulation causes a downward slope of isohalines towards the continental margin. If the mean flow at 1000 dbar is zero, Fig. 9 indicates eastward movement of the Atlantic water, in accord with Fig. 8. The sharp horizontal temperature gradient at the basin margin m Fig. 8 indicates that the rate of heat loss by upward diffusion from the Atlantic layer is larger over the continental slope than in the Canada Basra and that the cross-flow dtffuston of heat is relatively weak. Since there are indicattons that the advection speed over the slope is larger than in the basin, the cooling rate over the slope must actually be considerably enhanced. Over the 850-m nsobath north of Prudhoe Bay, meters deployed by the United States Coast Guard with the Frozen Sea Research Group in 1981 measured the 5-month mean flow at 270-m depth to be easterly at 2.5 cm s -j and occasionally speeds of 20 cm s -j were recorded (P. GREUSMAN, personal communication). From these data, nt appears that kinetic energy density is higher

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Pig. 9, Dynamic topography (dyn mm) relative to 1000 dbar o f the isobaric surface near the core of the Atlantic water layer (450 dbar). The downward slope of this surface to the north implies an eastward flow of Atlantic waters at 0.5 cm s-I relative to deep waters at 1000 dbar. Depth contours as in Fig. 8.

Oceanic thermal structure m the western Canadian Arctic

247

near the continental margin and could support a larger vertical d~ffuswity w~thm the Atlantic layer there. Over the deep shelves of the eastern Beaufort Sea a similar accelerated modification of Atlantic waters might be anucipated, although likely diminishing to the north as the increasing strength and compactness of the ice cover reduces the energy input to the ocean from the wind. Within the Atlantic layer m the western Archipelago, there is a significant loss of heat by diffusion, but only a minor loss of salt (Fig. 7). This reqmres either that the vertical diffusivities for heat and salt are equal while the curvature of the temperature profile exceeds that of the salinity profile, or that the vertical diffuswity for heat exceeds that for salt. The strong curvature in T-S relation at the core of the Atlantic water (~400 to 500 m; F~g. 2) ~s evidence for the occurrence of the former condition, while the latter circumstance is possible through double-diffusive convection. Since there have been observations within the Arctic thermochne of stepped profiles of temperature indicative of this phenomenon (NEAL et al., 1969) its significance in the Arctic Ocean thermal balance will be estimated. During the CTD programs m November 1979 and March 1981 steps in sahnity and temperature indicative of double-diffusive convection were observed in the southern Canada Basin over an area several hundred kilometers in extent. Those stations where stepped profiles were observed in 1981 are circled m Fig. 8. In 1981, observation of the steps during a second visit to one of the stations demonstrated that the phenomenon had hkely persisted for almost a month. A section from the stepped profiles which were observed m 1981 over a depth range of almost 150 m at Sta. 31 is shown in Fig. 10. Homogeneous layers whereto property transfer is convective are separated by layers with steep property gradients wherein transfer is diffusive (see TURNER, 1973, for discussion). The sharp transitions between the convectwe and the diffusive layers are smoothed to some degree because of the limited frequency and wavenumber bandwldths of the temperature and conductivity sensors on the CTD probe. However, since the thickness of the diffuswe layers (<0.5 m) and the time required to transit them (~<0.3 s) are considerably larger than the response distance (~0.1 m) and the response time (~0.05 s) of the sensors, the steepness of property gradients ts well represented. To illustrate this, sensor responses to idealized step inputs are also plotted in Fig. 10. The impulse response of this particular conductivity cell is based on TOPHAMand PERKIN(1984). Using appropriate values for the molecular diffuswlties (KT~ 1,5 × 10 -7 m 2 s-J; K s ~ 1.6 × 10 -9 m 2 s-J), the rates of transfer through the observed dlffuswe layers are ~0.02 W m -2 for heat and ~ 4 × 10-jj kg m -2 s -I for salt. Both these values are very small in comparison with estimates of basin-averaged fluxes near the top of the Atlantic layer, which are ~7 W m -2 and 5 × 10 -7 kg m -2 s J (AAGAARD and GREISMAN, 1975). Nonetheless they are m quantitative agreement with laboratory studies of double diffusive convection conducted at the values of the density ratio Rp = (f~AS)/(aAT) which are encountered in the Canada Basin (HuPPERT, 1971). Because of the limited range of temperature in the Arctic Ocean, 13/ct is approximately constant and Rp can vary significantly only if the slope of the T-S curve so varies. Above the temperature maximum in the Atlantic water, the largest value of ~T/OS encountered is such to reduce Rp only to about 7. The effective vertical diffuslvitles associated with these steps are only 3 to 10 times the molecular values for heat and salt. Double-diffuswe convection as observed in the lower thermocllne of the Arctic Ocean ~s apparently not capable of transporting heat upward at a significant rate. That the phenomenon is observed at all in the Canada Basin is indicative of the very small amount of kinetic energy available for turbulent mixing within waters overlying the abyssal plain. That the

248

H. M ELLINOel al.

0.665

CONDUCTIV[TY 0.66"/

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TEMPERRTuRE t .....

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....

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0.6"/1

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....

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Fig. 10. A subsection of hydrographic profiles observed at Sta. 31 (73°01'N, 134°50'W)during March 1981 (solid hnes). The stepped nature of the profiles is thought to result from double-diffusive convecuon which moves heat from the warm intrusive layer of water of Atlantic 6ngin into the cooler fresher water of the overlying Arcuc subsurface layer. Convectwe and dlffuswe components of the structure are ind=cated. The responses of the temperature and conducUvtty sensors to hypotheUcal steps having diffusive layers of zero thickness are plotted as dashed lines. The property gradients m the diffuswe layers are clearly not an artifact of fimte sensor response time.

phenomenon has not been observed, to date, in waters over the continental slopes or shelves suggests that there is considerably more kinetic energy available in these areas for turbulent mixing, a conclusion reached already on the basis of Fig. 8. Explanation of the observed modification of Atlantic waters (a cooling with negligible salt loss) near the continental margins is thus dependent upon the existence of a larger profile curvature for temperature than for salinity. On the basis of 20 profiles over a wide area of the southern Beaufort Sea in 1981 the mean ratio of ~ T / ~ z 2 to ~2S/~z2 is 6, and s.d. = 3. Thus if eddy diffusivities are equal for heat and salt, water-mass modification by diffusion will cause movement on a T - S diagram towards lower temperature with time along a line having a slope of 6. In Fig. 11, T - S diagrams for two station pairs over the continental slope in the Beaufort Sea are shown. Since the superimposed lines of slope 6 reproduce the direction of T - S evolution well, equal diffusivities for heat and salt in this area are implied. Both the baroclinic and barotropic flow components are potential sources of energy for the enhancement of vertical diffusivity within the Atlantic layer near the continental margin. Observations demonstrate that increases in geostrophic shear do occur over the continental slope in association both with shelf-edge upwelling in the wind-driven circulation (MELLING, 1983), and with the haline circulation which forms a density front near the shelf break. Geostrophic shears are, however, largely confined to the upper 200 m, far above the Atlantic temperature maximum at 450 m, and in the mean are associated with Richardson numbers ,,~100, although values below five are encountered within energetic mesoscale eddies. Therefore geostrophic shears cannot provide energy for mixing of Atlantic waters.

Oceanic thermal structure in the western Canadmn Arctic

249

SALINITY ~34

.'77 1

i

i

34 -~0 i

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13 14

407

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1103

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464

450

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CE ' £ECD L~J (3-

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Fig. 11. T-S curves for stations over the continental slope of the southern Beaufort Sea m 1981. Atlantic water temperature Js seen to decrease dramabcally over <30 km between Stas 103 and 104 and Stas 13 and 14. The arrow dehneates a line of slope 6 along whtch T-S values at the temperature maximum should progress if eddy dlffUStVmesfor heat and salt were equal Clearly the barotropic motions associated with tides or with wind forcmg of short period must provide the energy which supports diffusion within the Atlantic layer near the continental margin. Barotroptc motions wtll drive small-scale vertical mixmg only within the b o t t o m b o u n d a r y layer. Thus the observation that enhanced vertical diffustvity is found m precisely that zone o f the Beaufort Sea where the sea floor and the Atlantic temperature m a x i m u m coincide supports the deduction that mixing energy originates in the barotropic flow. Since tides in the C a n a d a Basin are very weak (£1 cm s - ' ; KOWALIK and MATHEWS, 1982) mixing over the contmentai slope by tides is hkely negligible, and occasional windinduced barotropic motions of the Atlantic waters must be largely responsible for upward mixmg. (Although no observations o f energetic internal tidal waves, which might break over the slope and cause mixing, exist, their contribution to mixmg cannot be ruled out.) A s uoted earher, US C o a s t G u a r d d a t a from 1981 north of A l a s k a show short-lived movements of Atlantic water at speeds as high as 20 cm s -t. Evidence that b o u n d a r y - l a y e r mtxmg does occur within the Atlantic water ts presented in Fig. 12 which depicts h y d r o g r a p h i c profiles from the continental slope of the Mackenzie shelf. A well-mixed b o u n d a r y layer 10 m in thickness is evident above the sea floor at 500 m. U s m g the estimate for vertical dtffusivity of ARMI (1979) and applymg it to the case o f Atlantic waters over the slope in the Beaufort Sea, where h -~ 5 m, x ,-, 5 d (the time scale o f meteorological events driving barotropic surges of current), dtffusivity is estimated to be h2/x ~ 6 × 10 -5 m 2 s -~. A s s u m i n g a balance between vertical diffusion and eastward advection at the Atlantic core in the southern Beaufort Sea (where 32T/az2,~ 2 × 10 -5 °C m - : and aT/ax ~ 0 . 0 1 3 o c per 100 km), an eastward flow u ~- 1 cm s -j is adequate to m a t c h temperature observattons (Ftg. 8) for this diffusivity. Smce this calculated value lies between the measured 5-month mean flow at 270 m over the slope

250

H. M ELLINGet al.

34 .S I

I

L

I

I

1

i

i

i

I

,

i

,

I

,

34 .9 I

I

35 .0 I

I

(OEG, C) 0.6 o.g

TEMPERRTURE 0.2 0.4

o0.0 L#~

SFILINITY 34 .7 34 .8

34 .6 i

,

I

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i

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l

.0 i

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24

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(73

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29,0

Fig 12 Hydrographfc profiles from 1981 over the continental slope of the southern Beaufort Sea The 10-m thick benthic boundary layer over the sea floor at 500 m is mdleatwe of significant barotroplc flow at the depth of the Atlantic water core (460 m) The water within the boundary layer originates at 800-m depth offshore

(--,2.5cms -~) and the baroelinic flow at 4 5 0 d b a r within the basin (,--0.5cms -j, re 1000 dbar), this diffusivity estimate is plausible. It exceeds the effective thermal diffusivity for the double-diffusive instability by a factor of 100. While the existence of a reservoir of cooled Atlantic waters on the western periphery of the Archipelago has been discussed, the mechanism driving this water over the sills bounding the large basms within it has not. The highest salinity o f water within the western Archipelago ~s a measure of the effectiveness of the sills along its western periphery in restricting mflow from the C a n a d a Basin. In recent years, the highest salinities observed (not necessarily the highest present) have been 34.88, 34.84, and 34.82 in western Amundsen Gulf (1977; MACDONALD et al., 1978), western Parry Channel (1982) and the Prince Gustaf Adolf Sea (1982), respectively. Water of these salinities ts found in the eastern Beaufort Sea at about 600, 430, and 400 m, respectively (1981 data). Upon reference to sill depths charted in Fig. 1, it is clear that water which origmates in the C a n a d a Basin at a depth greater than that of the entrance sill is present in at least the more southerly two of these basins. The required vertical displacement is about 240 m into Amundsen Gulf and 50 m into Parry Channel. However, since the 1981 observations in the Beaufort Sea indicate a downward tilting of these isohalines towards the continental margin, the mechanism leading to renewal of Archipelago deep waters must function either intermittently or on a small spatial scale. Within the sheltered secondary stlled basins in eastern Amundsen Gulf and M'Clintock Channel, deep-water sahnities o f 34.81 and

Oceanic thermal structure m the western Canadian Arctic

251

34.68, respectively, correspond to those at depths in the outer basins which are close to the depths of the intervening sills (335 and 290 m, respectively). The upward forcing of water over these sills in more sheltered waters is apparently weak. From the viewpoint of deep-water inflow over sills into the Archipelago, it is worthwhile noting that the water within the benthic layer at 500 m in Fig. 12 has the T - S properties of water found at 800 m in deeper waters nearby. Its vertical displacement against buoyancy forces ~s presumably a consequence of Ekman transport in the benthic boundary layer in response to a shoreward directed (barotropic at these depths) pressure gradient. At maximum displacement the upslope component of the pressure gradient force balances the downslope component of the buoyancy force. For the sea-floor layer in Fig. 12 this balance would occur for a barotropic pressure gradient sufficient to drwe a westward geostroph~c flow at 7 cm s ~. Thus modest barotropic pressure gradients established by meteorological systems could move deep waters over sills into the Archipelago through frictional effects. DISCUSSION

Water moving from the Arctic Ocean into the Archipelago within the halochne warms progressively with eastward and southward motion. While this warming is associated with a diffusive flux of heat from the underlying Atlantic layer, it is also dependent upon the absence in the Archipelago of the advective intrusions of near-freezing shelf waters which in the Arctic Ocean counteract the warming. Within the western part of the Canadian Arctic Archipelago, sea-ice conditions and seafloor topography are much different from those in the peripheral seas of the Arctic Ocean. Ice-cover is essentially fast for at least 6 months of the year, and as a consequence ice growth and salt rejection are limited by climate to about 2 m y-l and 50 kg m -2 y ~, respectively, even assuming the unlikely occurrence of completely ice-free waters at the onset of the freezing season (5/10 to 8/10 old ice coverage are typical except in Amundsen Gulf; MARKHAM, 1981). A salt flux of this magnitude into the uppermost 50 m of the water column is not adequate to enable significant penetration of surface waters into the pycnochne wherein salinity increases from 30 to 34. Moreover, the western basins of the Archipelago are deep (500 to 600 m) and steep-sided. As a result waters shallower than the 50-m mixing depth are very limited in area, and retention of large volumes of water near the freezing interface for the time reqmred for appreciable salinity changes is unhkely. For these reasons the production of cold source water for the Arctic subsurface layer is much less within the western part of the Arctic Archipelago than in the shallow shelf seas of the Arctic Ocean. The balance between intrusions of cold, saline (33 to 34) shelf water, the upward diffusion of heat and salt, and the export of pycnocline waters, which maintains the particular hydrography of the Arctic Ocean, is thus disturbed when waters enter the Canadian Archipelago. Without continuing cold water intrusions, the heat gained from the Atlantic layer within the Archipelago can have a noticeable warming effect on the pycnocline. Most of the heat gained by the halochne in the western Archipelago hes too deep to move directly over the sills (<100-m depth), limiting access into Lancaster Sound. It is, however, immediately accessible for mixing into the surface layer by the increasingly energetic flows m the shoaling and narrowing westward approaches to them. That such v~gorous upward mixing occurs at least in the northern approaches to Barrow Strait is demonstrated by the very linear T - S relationships (Fig. 2b) and by the numerous interior mixed layers (Fig. 13) which are observed m Byam Martin Channel and m Penny Strait. Flow kinetic energies are 10 to 100

252

H. MELLINGetal. SALINITY

SALINITY

31

32

I

i

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l

l

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33 ,

,

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. . . .

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. . . .

2'6 . . . . SIGNfl-T

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.

.

.

Fig. 13. Hydrographic profiles in Byam Martin Channel (Sta. C3) and Penny Strait (Sta. A3) m March to April 1982. The thin layers of uniform water property within the pycnocline are evidence for strong vemcal mixing m this area. Note that the near-surface temperature in Penny Strait is 0.1 °C above freezing temperature (' +'s).

times larger in these areas than in the deeper areas to the north and west (PECK, 1977, 1978, 1980a,c). The net effect of this mixing is a flux of sensible heat toward the surface. While in most places this flux is too small to prevent initial ice growth in winter, it is sufficiently large to slow further growth after an insulating layer of ice and snow has accumulated, and to promote early ablation of the ice cover, from below, when the flux of heat through the ice is reduced with increasing air temperatures in spring. The Arctic surface layer in March and April is generally observed to be isothermal within a few millidegrees of freezing over tens of meters. However, in Penny Strait, measurements in 1982 showed a gradient of temperature in the water below the ice and a temperature of 0.1 °C above freezing only 5 m from the surface (Fig. 13). Ice thickness was only 1 m, whereas a 2-m thickness at this season is typical in other areas. Because of the reduced ice thickness the flux of heat to the atmosphere at this site would be approximately twice that through the usual 2-m ice (,,-30W m-2), with the important difference that the heat lost from Penny Strait is entirely sensible heat from the water column (since the water is above freezing) rather than latent heat released by ice formation. In some places the turbulent transfer of sensible heat to the surface, and the turbulent agitation of any growing frazil ice are sufficient to prevent formation of a solid ice cover all winter. The polynyas which result are well-known recurrent features in Penny Strait and Queens Channel, in HelPs Gate and Cardigan Strait, and in Byam Martin Channel (STmuWG and

Oceanic thermal structure in the western Canadmn Arctic

253

CLEATOR, 1981). Detailed studies of air-sea mteractlon at the Dundas Island polynya (TOPHAM et al., 1983) have demonstrated that the flux of heat to the atmosphere from open water can exceed 500 W m -2, and that a sizeable fraction of this flux originates as sensible heat m the water column, since surface water temperature is typically 0.1 to 0.2°C above freezing. In western Lancaster Sound, observations of a cold haloclme sublayer have frequently been taken as evidence for the origin of this sublayer in the Arctic Ocean (MUENCH, 1971). However, the depths of the limiting sills (,,,100 m) are considerably shallower than the bottom of this layer in Lancaster Sound, and the halocline upstream of the sills is considerably warmer than its Arctic Ocean precursor. The observations made in western Lancaster Sound in March 1982 reveal intrusions of colder water at depths up to 300 m (Fig. 14). These data indicate that there is an active regeneration in this area of a cold pycnocline such as is characteristic of the Arctic Ocean, but not of the western Archipelago. The low temperature of the intrustve waters requires an origin at the ice-water interface. However, since the range m surface salinity of potential source waters m Barrow Strait and Penny Strait in March 1982 (32.2 to 33.0; Fig. 15) would enable descent only to 75 m in Lancaster Sound, the higher surface salinities required for descent to ,,-200 m must be produced nearby, most probably by freezmg over open water beyond the edge of fast ice in Lancaster Sound. This edge lay 3 to 6 km east of the observational sites in 1982. The very limited wintertime observations east of

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Fig.

the fast-ice edge in Lancaster Sound suggest that the surface convective layer may extend to 200 m or more in this area (LEMON and FISSEL, 1982). Thus the deep cold intrusions observed in March 1982 could well be indicative of interleaving with the waters of such a deep surface mixed layer existing further east. In M'Clure Strait the T - S characteristics of the inflowing water in the halocline point to a source in the eastern Beaufort Sea but the origin of tts m o m e n t u m is not established. Current measurements over the outer shelf of the southern Beaufort Sea have suggested the existence o f an eastward flowing undercurrent within the halocline over both the north Alaskan and M a c k e n z m shelves (AAGAARD,1983; HUGGETT et al., 1977; Frozen Sea Research Group, unpublished data). However, since the deep (,-,360 m) entrance to Amundsen Gulf presents an effective barrier to a northward continuation of this current west of Banks Island, the probable region of origin of the undercurrent in Parry Channel is the continental slope west of Banks Island. Little of the oceanography of the area to the west of Banks Island is known. However, observations of ice drift dtscussed by MELUNG and LEWIS (1982) and by STIRLING and CLEATOR (1981) do show it to be an area of persistent offshore Ekman transport and southward wind stress. Conditions are thus favourable for the formation of a northward winddriven undercurrent near the shelf break such as observed on the Pacific coast of North America (HIcKEY, 1979) and modelled by MCCREARY (1981). Local bathymetry near M'Clure Strait could steer such an undercurrent eastward along the north coast of Banks Island.

Oceamc thermal structure m the western Canadian Arctic

255

Within the Canada Basin, the zone over the continental slope of the southern Beaufort Sea in which Atlantic water heat loss is enhanced is that same zone into which currents of shelf water at freezing temperature interleave at lesser depths (MELLING and LEWIS, 1982). The upward flux of heat from the Atlantic layer is thus trapped below ~ 100 m by the presence of a substantial heat sink and is not capable of influencing the surface heat balance and sea-Ice cover. Beyond the continental slope flows are at times sufficiently quiescent that doublediffusive mstabiliUes may develop. The effective thermal diffusivity associated with these instabilities is only ,-,0.01 times that characteristic of the slope region. Vertical exchange processes occurring along the periphery of the Arctic Basins are thus clearly of crucial importance in the thermal and mass balances in the upper 1000 m of the entire Arctic Ocean. CONCLUSIONS

Because the amount of seasonal ice growth and the area of shallow (,,~50 m) shelves are limited, cold brines produced through lee growth are not an important influence on the halocline m the western Canadian Archipelago (western Parry Channel, waterways of the Queen Elizabeth Islands, Lincoln Sea). As waters from the Arctic Ocean move east and south into the Archipelago, those within the halochne warm as a result of a diffusion of heat from underlying waters of Atlantic origin. In Parry Channel warming is greatest (0.25°C) at a salinity of 33.5. In western Parry Channel, the slowly decelerating inflow within the halocline may originate as a wind-driven undercurrent over the shelf west of Banks Island. The halocline inflow to Amundsen Gulf may be forced similarly or perhaps be a consequence of thermohaline flows on the Mackenzie shelf to the west. Forcing of inflows to the Queen Elizabeth Islands and to the Lincoln Sea is not clear. The existence of polynyas in the vicinity of shallow ( ~ 100 m) sills m Penny Strait, Hell's Gate, and Byam Austin Channel is directly attributable to the existence of sensible heat within the halocline where it is available for upward mixing in the strong turbulent flows m these constricted channels. The estimated loss of sensible heat to the atmosphere in these areas is a sizeable fraction of the total heat loss, and as a consequence the growth of an Ice cover is either inhibited or completely prevented. Without the steady warming of the halocline in the areas upstream of these channels, sensible heat would remain too deep m the water column to be effectively mixed to the surface in these areas. In winter, a near-freezing pycnochne, such as found in the Arctic Ocean, is regenerated in western Lancaster Sound both through the descent of surface waters of higher salinity confluent from the north and west and through the effects of deep vertical convection associated w~th ice growth on open water east of the fast-ice edge. Waters of Atlantic origin in all the basins of the western Canadian Archipelago are conSlstently cooler than those of the same salinity in the Canadian Basin. The d~fference increases with salinity to N0.1 °C m the deepest parts of the basins. In the western Parry Channel, the T - S characteristics of the waters below sill depth have remamed unchanged, within observational error, over 30 years. This implies a continual replemshment of these deep waters from a reservoir which has the same T - S characteristics. Deep-water renewal may be a result of upslope Ekman transport in response to barotropic forcing, There is evidence to demonstrate that this reservoir of cool Atlantic water occupies the regions on the margins of the Canada Basin where water depth is comparable to the depth

256

H. MELLINGet al. t

o f the temperature m a x i m u m in the Atlantic layer. In the southern C a n a d a Basin this correspondence o f depths occurs in a narrow zone over the continental slope, wherea~ on the eastern periphery o f the Basin it occurs in a broader band over the continental shelf. In western P a r r y Channel steady Atlantic water-mass characteristics can be maintained if the upwelling caused by deep-water renewal is counteracted by vertical diffusion. Balance can be achieved for both heat and salt at an uplifting speed o f 15 c m d - l , or a corresponding deepwater inflow at a rate of 0.1 x 106 m 3 s -l. The cooling o f Atlantic waters in the C a n a d a Basin which produces the modified water m a s s found over the continental margins occurs with little reduction in the salinity o f the waters involved. This circumstance is a consequence o f the larger curvature o f the profile o f temperature, rather than a reflection o f a large difference in the eddy diffusivities for heat and salt. Double-diffusive convection is not a rare phenomenon in the C a n a d a Basin where it occurs in the lower thermocline between 250 and 450-m depth. The fluxes o f heat and salt transferred u p w a r d by this mechanism are, however, very small (,,,0.02 W m-Z; ,,,4 x 10 T M kg m -2 s - l ) relative to averages over the Arctic Basin o f fluxes from the Atlantic layer ( ~ 7 W m-E; -,,5 x 10 -7 kg m -2 s-I). H e a t loss by the Atlantic water in the Beaufort Sea is more rapid over the continental margin than further offshore. The coincidence o f sea floor and Atlantic water depths in this zone suggests that the enhanced diffusivity is a result of turbulence in the benthic bot, ndary layer associated with the b a r o t r o p i c flow. Atlantic waters near 450-m depth in the southern Beaufort Sea move from west to east. D y n a m i c calculations suggest the average speed in the C a n a d a Basin is <0.5 c m s -~ . Over the continental slope, direct measurements indicate a larger average speed o f 2.5 c m s -l, and surges up to 20 c m s -i are encountered. In some respects, the western part o f the C a n a d i a n A r c t i c Archipelago can be regarded as a l a b o r a t o r y in which changes in waters which originate in the pycnocline o f the Arctic Ocean can be observed in the absence o f the intrusions o f near-freezing water which condition the pycnocline o f the Arctic Ocean. Within the Archipelago, because o f the continued upward diffusion o f heat from the Atlantic layer, the halocline w a r m s with e a s t w a r d movement. The existence o f p o l y n y a s and thinner seasonal ice in the central Archtpelago is a graphic illustratton o f the effect that Atlantic water heat could have on the ice cover o f the Arctic Ocean itself were it not for the d o w n w a r d flux o f cold water from surrounding shelves. This cold water ts a subsurface sink for upward diffusing heat which thereby maintains small temperature gradients and heat fluxes m the upper 150 m. Its intrusion into the pycnocline limits the depth o f surface mixing by forcing upwelling to counteract entrainment at the base o f the surface mixed layer.

Acknowledgements--These studies were supported by the Canadian Department of Fisheries and Oceans and by

the Canadian Department of Transport through the Energy Transport Research and Development Program. Logis0cal support and aircraft were provided through the Polar Continental Shelf Project of the Canadian Department of Energy, Mines and Resources. We thank those people wRhin the Polar Continental Shelf Project, the Canadian Hydrographic Service, Panarctac Ods Ltd., Bradley Air Servtces Ltd., Kenn Borek Air Ltd., and Quasar Helicopters Ltd. who contributed to the success of the field studies. We are indebted to our colleagues in the Frozen Sea Research Group and in Arctac Sctences Ltd. who were revolved in the many and vaned tasks associated with such programs, and to Andrew Wharton of Interact CompuUng Services Ltd. for his patience and competence m data processing. Constructwe cntac~sm by Paul Gre~sman of Dobrocky Seatech Ltd. ts greatly appreciated.

Oceanic thermal structure in the western Canadian Arctic

257

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