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Physical control of spring–summer phytoplankton dynamics in the North Water, April–July 1998 ! Tremblayc, B. LeBlanca, Z.-P. Meia,*, L. Legendrea,1, Y. Grattonb, J.-E. C.J. Mundyd, B. Kleina, M. Gosseline, P. Larouchef, T.N. Papakyriakoud, C. Lovejoya, C.H. von Quillfeldtg a
D!epartement de biologie, Universit!e Laval, Qu!e., QC G1K 7P4, Canada b INRS-Eau, 2800 rue Einstein, Sainte-Foy, QC G1K 4C7, Canada c Department of Biology, McGill University, 1205 Dr. Penfield, Montreal, QC H3A 1B1 Canada d Centre for Earth Observation Science, Department of Geography, University of Manitoba, Winnipeg, MB R3T 2N2, Canada e Institut des sciences de la mer (ISMER), Universit!e du Qu!ebec a" Rimouski, 310 All!ee des Ursulines, Rimouski, QC G5L 3A1, Canada f Institut Maurice-Lamontagne, Minist"ere des P#eches et des Oc!eans, C.P. 1000, Mont-Joli, QC G5H 3Z4, Canada g Norwegian Polar Institute, Troms, Norway Received 6 November 2000; received in revised form 23 July 2001; accepted 31 October 2001
Abstract A 4-month multidisciplinary expedition, beginning at the end of winter to track the spring phytoplankton bloom to its termination in summer, was conducted from April to July 1998. The aim of the expedition was to investigate possible mechanisms responsible for the high biological productivity of the North Water, the most productive and largest polynya in the Northern Hemisphere. The aims of the present study were to investigate: (1) the effects of the physical forces, driving the formation of the polynya, on the dynamics of the phytoplankton stock in the polynya over the spring–summer period; and (2) the factors that limit the maximum biomass of phytoplankton. Contrasting physical characteristics, including ice concentration, surface mixed-layer depth (MLD), salinity and temperature in the surface mixed layer (SML), were observed between the east and west sides of the polynya. The Greenland (eastern) side of the polynya was characterized by a shallow SML, warm temperature, and high salinity relative to the Ellesmere Island, Canada (western) side. Chlorophyll a (Chl)>1 mg m3 was observed in late April on the eastern side, and in late May on the western side. The peak phytoplankton bloom occurred in the southeastern part of the polynya, with average Chl of 15 mg m3 (240–300 mg m2) in the euphotic zone during the end of May and beginning of June. The increased phytoplankton biomass was associated with higher salinity and warmer temperature on the eastern side of the polynya. Low temperature in April and May decoupled the increase of Chl biomass from the shallowed SML, as predicted by Sverdrup’s model. As Chl in the euphotic zone increased to 5 mg m3, the proportion of light absorption by phytoplankton could not increase further with Chl biomass, which might have limited the increase of primary production in the water column. Although the initial nutrient inventories largely determined the maximum biomass of phytoplankton, self-shading occurred in the build-up phytoplankton biomass to B5 mg m3, which retarded the timing of the peak bloom. Both sensible heat due to deep warm water entrainment into the SML and the biological heating
*Corresponding author. Tel.: +1-418-656-7777; fax: +1-418-656-2339. E-mail addresses:
[email protected] (Z.-P. Mei). 1 Laboratoire d’Oceanographie de Villefranche (LOV), BP 28, 06234 Villefranche-sur-Mer Cedex, France. 0967-0645/02/$ - see front matter r 2002 Elsevier Science Ltd. All rights reserved. PII: S 0 9 6 7 - 0 6 4 5 ( 0 2 ) 0 0 1 7 3 - X
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effect via phytoplankton light absorption appear to contribute to the pattern of phytoplankton distribution in the North Water. r 2002 Elsevier Science Ltd. All rights reserved.
1. Introduction Polynyas are areas of partially or totally ice-free waters at times when surrounding waters are icecovered. The North Water is one of the largest recurring polynyas in the Northern Hemisphere. It occupies Smith Sound and the northern part of Baffin Bay, between Greenland and Ellesmere Island, and is a well known feeding and breeding ground for large populations of marine mammals and birds. Two mechanisms for the opening of polynyas have been proposed: latent heat and sensible heat (Smith et al., 1990). In the first case, the polynya is kept open by wind-driven advection of ice; the heat required to balance the loss to the atmosphere is provided by the latent heat from ice fusion as the ice continuously forms because of low air temperature. In the second case, the heat required to maintain the polynya open originates from oceanic sources, e.g., the upwelling of deep warm water (Smith et al., 1990). Latent heat polynyas are subject to deep mixing due to input of wind energy at sea surface and salt ejection during new ice formation. In contrast, sensible heat polynyas are subject to stabilization of the upper water column due to relatively warmer water and subsequent ice melt. According to the Sverdrup (1953) model, initiation of the spring phytoplankton bloom results from a combination of factors that include irradiance at the sea surface, which changes significantly from high to low latitude (Campbell and Aarup, 1989), attenuation of light in the water column by particulate and dissolved organic matter, which is a function of trophic conditions (Kirk, 1994), and thickness of the surface mixed layer, which determines the average irradiance to which phytoplankton are exposed. The original model was challenged in specific waters by Townsend et al. (1992) and Eilertsen (1993), and in polar seas by Nelson and Smith (1991) and Boyd et al. (1995). Phytoplankton development in polynyas,
formed by different heat mechanisms, may be influenced by the timing of polynya opening, ice cover, and water-column mixing; the relative importance of these factors is little explored. An earlier study showed latent heat to be mostly responsible for the opening of the North Water (Muench, 1971). During a 2-day cruise in May 1991, Lewis et al. (1996) inferred an input of sensible heat from depth to the upper layer in the region, to which they linked the very high phytoplankton biomasses they observed. The first objective of the present study was to examine the hypothesis that variations in phytoplankton biomass in the North Water are related to spatiotemporal changes in the physical characteristics of the upper water column, with reference to different heat inputs. The second objective was to investigate the potential maximum phytoplankton biomass and its physical and chemical constraints within the polynya. Moderate to high phytoplankton biomass or productivity have been reported in some polynyas, e.g., the Northeast Water (Smith, 1995; Pesant et al., 1996) and the North Water (Lewis et al., 1996) in the Arctic and the Ross Sea Polynya in the Antarctic (Arrigo and Weisse, 1998). The timing and yield of phytoplankton blooms in marine waters determine the coupling between primary production and the heterotrophic food web and, hence, the fate of primary production (Legendre, 1990). The timing of bloom initiation and seasonal variation in the stock of the North Water were poorly known. Recent studies on the coupling between physical and biological processes in polar seas, including polynyas, have shown that phytoplankton biomass, composition and production are highly dependent on the hydrographic characteristics of the coastal, shelf or open-water region (Arrigo et al., 1999; Pre! zelin et al., 2000; Smith and Asper, 2001). Phytoplankton growth causes changes in the characteristics of the water column, both
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physical (e.g. underwater irradiance and stability of the upper water column due to the biological heating effect; Stramska and Dickey, 1993; Morel and Antoine, 1994) and chemical (e.g. nutrient concentrations and the partial pressure of gases; Yager et al., 1995; Arrigo et al., 1999). Phytoplankton not only actively adapt to a specific environment, but also modify biogeochemical cycling in the marine ecosystems they support (Verity and Smetacek, 1996). Hence, the interactions between phytoplankton and physical processes in the upper water column have direct consequences on atmosphere–ocean interactions, with respect to exchanges of heat and gases.
2. Materials and methods 2.1. Sampling Sampling was conducted in the North Water on board the Canadian icebreaker Pierre Radisson from 7 April to 20 July 1998. Because sampling was limited to the western half of the polynya until 21 April, the data presented here were from 21 April through 20 July (Fig. 1). Water was collected with a General Oceanic rosette sampler equipped with a CTD profiler (Falmouth Scientific Inc ICTD), a Seatech fluorometer and 24, 10-l bottles (Brooke Ocean Technology Limited, Dartmouth, Nova Scotia) at six optical depths in the euphotic zone (nominally 100%, 50%, 30%, 20%, 10% and 1% of surface irradiance) and several depths below. At 30% of all stations (‘‘full’’ stations) the euphotic zone, from surface to the depth of 1% surface photosynthetically active radiation (PAR, 400–700 nm), was determined as described below. For the remaining stations, the euphotic zone was corrected a posteriori based on the polynomial relationship between measured euphotic zone depth (Y; m) and chlorophyll a (Chl) concentration (X, average Chl in the euphotic zone as sampled; mg m3): Y ¼ 0:0412 X 3 þ 1:1921 X 2 211:136 X þ 55:024 (n ¼ 48; r2 ¼ 0:82). Note that this equation was used only to obtain the integration depth for calculating Chl in the euphotic zone. Water samples for the determina-
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tion of Chl were prefiltered through a Nitex screen (300 mm in April, and 500 mm in May–July). 2.2. Ice concentration Ice-concentration data were extracted from a 2 2 km2 grid, which contained the percent ice cover derived from classified RADARSAT images (Mundy, 2000). The classification distinguished among various types of ice (new/gray ice, gray– white floes, medium first-year ice, thick first-year ice, multi-year ice, rough/rubble ice, and brash ice). The percent cover of each ice type was averaged over the 10 10 km2 area surrounding each station. The modal thickness of each ice type was assigned as follows: new/gray ice, 15 cm; gray– white floe, 30 cm; medium first-year ice, 100 cm; thick first-year ice, 200 cm; multi-year ice, 400 cm; rough ice, 400 cm; and brash ice, 75 cm. The weighted average thickness of ice at each station was calculated from the average thickness of each type of ice and its percent coverage averaged over three successive days (the day of sampling and the two preceding days). This approach provided a first-order approximation of ice concentration around the stations where biological samples were taken. Detailed information on the spatial patterns of ice cover and ice motion in the polynya can be found in Mundy and Barber (2001) and Wilson et al. (2001), respectively. 2.3. Temperature, salinity and mixed-layer depth (MLD) Temperature and salinity were recorded over the water column and interpolated at 1-db intervals (ca. 1 m). Mean temperature and salinity were calculated with the interpolated data by integration between the shallowest measurement and the bottom of the surface mixed layer (SML), with the shallowest valid data point on the profile taken as the surface value. The surface MLD was defined as the depth where potential density was 0.01 higher than at the shallowest measurement depth or 0.04 when there was abnormal fluctuation of the density profile in the SML, according to Schneider and Muller (1990). Vertical stability of the water column was characterized by the Brunt–V.ais.al.a
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Fig. 1. Map showing maximum area of the North Water, with positions of sampling stations for hydrographic and biological samples. Numbered stations were sampled from April to June: lettered stations (D1, E1, E2, N1, N2 S1, S2, S4, and S5, the sites for current meter and/or sediment trap moorings), in July.
frequency (N 2 ; 104 s2; Pond and Pickard, 1983). When multiple CTD casts were available at a station and the inter-cast variation was small, the average MLD for the station was calculated. When the inter-cast MLD varied widely, N 2 and in situ fluorescence profiles were compared; the N 2 profiles most relevant to fluorescence were used to determine the MLD. 2.4. Photosynthetically available radiation In ice-covered areas, the coefficient of diffuse attenuation (kice ; m1) of the photosynthetically
available radiation (PAR) by the ice (including the snow cover and ice algae) was estimated at several ice stations. Coefficient kice was calculated with the exponential attenuation model (e.g. Arrigo et al., 1991), using the percentage of sub-ice PAR (measured with a LI-COR 190SA underwater quantum sensor) relative to PAR above ice (LICOR 192SA quantum sensor) and the ice thickness measured in situ at several ice stations (Nozais et al., 2001). Since there were no clear relationships between kice and different types of ice with different thicknesses, an average kice value for all of the in situ measured stations was used for
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each month; i.e. 4.3671.77, 3.8871.44, and 1.8270.95 m1 for April, May, and June, respectively. For each ice type, the percent sub-ice PAR was obtained from surface irradiance, the average thickness of the type (see above) and the corresponding coefficient of diffuse PAR attenuation. Incident PAR was measured with an on-deck PAR sensor (LI-COR LI-190 SA cosine corrected flat sensor) during the 4-month sampling period. At stations where ice was present, a weighted average sub-ice PAR was calculated based on the percent cover of the various ice types and their thickness, over an area of 10 10 km2 around each sampling station (see above). Even though this approach provides simplified estimates of light penetration through the ice, assuming an even distribution of snow on top of the ice, a comparison of the estimates with values measured in situ at ice stations showed close correspondence between the two sets of values. A detailed model of light transmission through snow and different types of ice is under development (D. Barber, pers. comm.). At sampling stations with open-water, sea-surface irradiance was calculated from incident ondeck PAR by using a surface reflectance of 7.5%. When ice was present, the sub-ice PAR (previous paragraph) of the ice-covered area was added to the open-water value (sea-surface irradiance) to obtain the corrected sea-surface irradiance at the station. An underwater PAR sensor (Satlantic SPMR with 13 visible-light channels) was used to determine PAR profiles at full stations (ca. 30% of the sampled stations). The attenuation coefficient of downwelling PAR (kd ; m1) in the water column was interpolated at 1-m intervals. The interpolated values were used to calculate average kd in both the euphotic zone and the SML, taking into account the uneven vertical distribution of particulate matter in the water column. Mean PAR down to depth z was calculated as E¼
E0 ½1 expðkd zÞ ; kd z
ð1Þ
where E0 is the daily averaged PAR at surface (Riley, 1957, 1967). In order to smooth out the
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large day-to-day variations in incident PAR, daily averaged PAR at each sampling station was calculated by averaging over three successive days, i.e. the day of sampling and the two preceding days (Levasseur et al., 1984). As the ship was moving during the 3-d sampling, the 3-d averaged PAR represents the average incident irradiance over the sampling area. The cloud conditions and incident PAR were assumed to be uniform over the sampling area.
2.5. Nutrients and chlorophyll a Concentrations of dissolved inorganic nitrogen (DIN; nitrate+nitrite) were determined on board the ship using an ALPKEM autoanalyzer, following routine colorimetric protocols (Grasshoff, 1976). For Chl, Nitex-prefiltered water samples were filtered onto Whatman GF/F (nominal pore size of 0.7 mm, total Chl) or 5-mm pore-sized Nuclepore polycarbonate filters (large Chl, >5 mm), followed by acetone (90%) extraction of the filters in a cool (B51C) dark environment for 24 h, before fluorometric determination. The Chl concentration was corrected for phaeopigments by acidification of the extract (JGOFS, 1996). The mean concentrations of Chl (mg m3) and nutrients (mmol m3) in the euphotic zone or the SML were calculated from vertically integrated values (mg m2) between surface and bottom of the euphotic zone or the SML (or above, when data were not available at the bottom of the SML), divided by the respective depths (m). Phytoplankton samples collected at the depths of 50% and 1% of surface PAR during the full stations (ca. 30% of all stations) were enum. erated using Fluorescence Nomarski Utermohl (FNU) microscopy, as described in Lovejoy et al. (1993). In the present study, the organisms were stained with Hoechst 33342 and examined at a magnification of 400X. Samples from areas with low Chl (o1.5 mg m3) were concentrated as in Lovejoy et al. (2000). Although we counted cells o5 mm, that part of the community was probably underestimated under the magnification used.
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2.6. Statistical analyses Model II linear regression analysis was used for determining the relationships of Chl with DIN (standard major axis procedure, SMA, or reduced major axis, RMA) and of kd with Chl/ Chl+phaeopigments (ordinary least square, OLS, for predictive purposes) (Sokal and Rohlf, 1995; Legendre and Legendre, 1998). The relationship between the proportion of light absorbed by Chl+phaeopigments in the euphotic zone (A, %) and Chl+phaeopigments was determined using the model of Webb et al. (1974), as suggested by Frenette et al. (1993) but forcing the fitted curve to pass through the upper bound of the data points. The relationship between the depth of the euphotic zone and Chl was fitted with an exponential decaying model based only on data obtained at stations where in situ measured euphotic zone depths were available. In order to identify the major physical variables that could explain the abundance of Chl at the different sites, the stations were divided into three groups based on the average total Chl in the SML, followed by multiple discriminant analysis (MDA; Legendre and Legendre, 1998) of the three groups by physical variables (MLD, average N 2 of the SML, average N 2 over the pycnocline, average temperature, and salinity of the SML). Computations were carried out using the Statistica software (StatSoft, Tulsa, OK). Raw data were standardized before the pooled-within-groups correlations were calculated, as recently suggested by Tabachnick and Fidell (1996).
3. Results 3.1. Ice concentration Spatial distributions of ice concentration expressed as average ice thickness over 10 10 km2 areas around each sampling station (Fig. 2) showed that in late April, the average ice thickness ranged from 60 to 130 cm. Near Greenland, the ice thickness was 60–100 cm, except at station 33 where it was 29 cm, compared to 100–130 cm near Ellesmere Island. The ice thickness decreased
drastically in May, with little ice on the eastern side except close to the coast, and values on the western side ranging from 20 to 130 cm. The southernmost transect (75.251N) was still covered with ice then and was only visited in June when ice was disappearing from the polynya. During June and July, most of the polynya was free of ice (data not shown), except for episodic floes moving southwards. 3.2. Spatial distributions of chlorophyll a Chl values were lowest in late April, when the mean concentrations over the euphotic zone were low (o2 mg m3; Fig. 3) and the highest values (>1 mg m3; 50–70 mg m2) occurred at the eastern end of transects (Fig. 3). The main increase in Chl occurred during May, when there was a generally increasing gradient of Chl from north to south and, on each transect from west to east. The highest values, up to 15 mg m3 (240 mg m2), were observed in the southeast of the polynya, while values ranging from 0.1 to o 5 mg m3 (o120 mg m2, most of them o60 mg m2) were observed in the north and west. The southernmost transect (75.251N; outside the polynya) was not sampled in April and May, because of heavy ice cover. In June, the mean Chl concentration of 6.49 mg m3 (158 mg m2) was slightly higher than in May (4.24 mg m3; 101 mg m2); the June value is an average for all the sampled stations except those on the southernmost transect (values o1 mg m3; o60 mg m2). Chl decreased, however, to o2–6.7 mg m3 (p143 mg m2) in the southeast of the polynya, where it had been the highest in May. In July, Chl declined to ca. 1– 6 mg m3 (o151 mg m2). Hence, high Chl biomass in the polynya persisted over 1.5 months, from late April, when values >1 mg m3 (>50 mg m2) were observed in the southeast of the polynya (Fig. 3). From April through July, the horizontal distributions of >5 mm Chl (% of total Chl; Fig. 4) were similar to the overall patterns of total Chl (Fig. 3). In April, when total phytoplankton biomass was low, large phytoplankton accounted for 57% (range of 35–80%) of total Chl, with the highest values (>70%) occurring on the
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Fig. 2. Horizontal distribution of ice thickness (cm) in the North Water during April and May, where average values were calculated over 3 d (day of sampling and the two previous days) for a 10 10 km2 grid around each station. Dots indicate positions where satellite-image data were obtained.
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Fig. 3. Horizontal distribution of average total Chl (mg m3) in the euphotic zone during April, May, June, and July; note different scale for April. Dots indicate positions of sampling stations.
eastern side (e.g. stations 18, 36, 38 and 40). Two stations on the western side also showed high values (68% and 79% for stations 44 and 47,
respectively). In May, the >5 mm Chl increased to 78% of total Chl (range of 55–100%), showing that large phytoplankton were mostly responsible
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Fig. 4. Horizontal distribution of the fraction of >5 mm to total Chl (expressed as %) in the euphotic zone during April, May, June, and July. Dots indicate positions of sampling stations.
for the bloom. In June, the >5 mm Chl continued to dominate, with an average proportion of 86% (70–100%). In July, the >5 mm Chl started to
decrease together with total Chl, the large phytoplankton then accounting for 72% of total Chl (31–87%).
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Fig. 5. Scatter plot of the contribution of diatoms to the total number of >5-mm autotrophic phytoplankton (expressed as %) as a function of Chl concentration (mg m3) in the samples (pooled from 50% and 1% surface irradiance).
Diatoms numerically dominated the >5-mm phytoplankton community. The contribution of diatoms to phytoplankton cells >5 mm (pooled data from the 50% and 1% photic depths) increased in parallel to Chl. When Chl reached 3.5–5.0 mg m3, the diatoms accounted for 90– 100% of the >5-mm phytoplankton community (Fig. 5). 3.3. Chlorophyll a, mixed-layer depth and water masses In late April, Chl >1 mg m3 (Fig. 6a) was associated with the water mass with the highest salinity (>33.2). The water mass with salinity o33.0 was characterized by low Chl (o0.5 mg m3). The surface water temperature was at the freezing point (o 1.71C) throughout the polynya due to low air temperature. The water with salinity >33.2 is believed to have originated from Baffin Bay Water, while lower-salinity water (32.0–33.2) originated from the Arctic outflow or resulted from mixing of the Arctic outflow with Baffin Bay water (Tremblay et al., 2002; B#acle et al., 2002). In May, a major phytoplankton bloom developed in the southeast of the polynya,
where salinity was >33.0 and temperature ranged from 1.41C to 0.311C, whereas Chl remained low (o3.5 mg m3) at stations with temperature close to freezing (o1.41C) and lower salinity (o33.0). In June, the highest Chl occurred in water with low salinity (as low as ca. 32.5) and higher temperature (from 11C to 11C) than in May. In July, moderately high Chl (5–7 mg m3) occurred at four stations with salinity o31.5 and temperature ranging from 1.51C to 1.01C. In contrast to the deep SML of 60–105 m observed at stations with low salinity (o33), shallow SML of B20 m at stations with high salinity (33.2–33.5) and Chl (1.2–2.3 mg m3) were observed in late April (Fig. 6b). In May, the highsalinity (>33) stations had a shallow SML (9– 25 m) and temperature >1.51C, whereas the SML was deep at stations with salinity o33 and temperature o1.51 C. The only exception was station 18 with salinity of 33.5 and temperature of 1.41C, but with MLD of 59 m. In June and July, the SML was shallow at most stations, ranging from 5 to 25 m, with a mean MLD of 12 m, irrespective of salinity and temperature. The highest MLD values were observed at northernmost stations in June.
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Fig. 6. Bubble plots of (a) average Chl (mg m3) in the SML and (b) MLD (m) on the T–S diagram, where bubble diameters are linearly proportional to the values.
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3.4. Effects of physical characteristics of the water column on total chlorophyll a
kd ¼ 0:090 þ 0:017½Chl þ Phaeopigments
Multiple discriminant analysis (MDA) was used to investigate how total Chl was controlled by physical variables (i.e. MLD, average N 2 in the SML and the pycnocline, average salinity and temperature in the SML) in April, May, and June (Fig. 7). In order to do so, stations sampled in late April, May, and June were first divided into three groups (group members were different for the 3 months) based on average Chl in the SML (Table 1). In June, three stations (2, 4 and 7) at the northern tip of the polynya were not included in the analysis, because of distinguishable hydrographic properties from the rest of the polynya, judging from CTD profiles (B#acle et al., 2002) and taxonomic data (Lovejoy et al., 2002). In July, there were not enough sampled stations to conduct valid MDA. For the three sampling months, the three groups were mostly discriminated on the first canonical axis, which explained, from 92% to 98% of the total variance. The first discriminant function was mostly correlated with the following variables (Table 2 and Fig. 7): salinity and MLD, in late April; mostly temperature, followed by salinity, MLD and the average stability of the SML, in May; and temperature, followed by the stability of the pycnocline, in June.
In Eqs. (2) and (3), the slope represents the attenuation due to phytoplankton pigments, while the intercept provides an estimate of the attenuation due to components other than phytoplankton pigments (i.e. seawater, dissolved organic matter and particulate detritus, kw ). The proportion of light attenuated by the pigments (A, %) was calculated as
ðn ¼ 48; r2 ¼ 0:80Þ:
A ¼ ðEw EÞ=Ew ;
The kd for PAR was significantly related to either the mean concentrations of Chl or Chl+phaeopigments in the euphotic zone. The corresponding OLS regression equations are: kd ¼ 0:094 þ 0:017 ½Chl
ðn ¼ 48; r2 ¼ 0:79Þ; ð2Þ
ð4Þ
where Ew is the mean irradiance in the euphotic zone without attenuation by phytoplankton pigments. Ew was calculated using Eq. (1), with kd in the equation being replaced by kw (given that in situ kd includes the attenuation by both phytoplankton pigments and non-phytoplankton components in seawater). There was a strong exponential relationship between A and mean Chl+phaeopigments in the euphotic zone (Fig. 8a). The upper bound of the scatter plot was fitted with a model similar to the photosynthesis-irradiance relationship of Webb et al. (1974), replacing here the term for the saturated photosynthetic rate (Pmax ) by Amax ¼ 62%; and E by Chl+phaeopigments: A ¼ 62f1 exp½39ðChl þ phaeopigmentsÞ=62g ðr2 ¼ 0:73; n ¼ 49Þ:
3.5. Chlorophyll a and irradiance
ð3Þ
ð5Þ
The upper bound A ¼ 62% was approached when Chl+phaeopigments >5 mg m3. The depth of the euphotic zone (zeu ) declined exponentially with increasing Chl (Fig. 8b). Note that this analysis only included the stations where both Chl and the depth of the euphotic zone were measured in situ. For Chl >5 mg m3, zeu leveled off at 16–25 m (asymptotes of two exponential negative functions for the lower and upper bounds
–––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––––" Fig. 7. Discriminant analyses of station groups based on average Chl (mg m3) in the SML using physical variables (see Table 1) for April, May, and June. Closed lines provide a visual reference for the groups, against which discriminant analysis was conducted. In April, Axis I was most correlated (in the positive direction) with salinity; Axis 2, with N 2 of the pycnocline, followed by average N 2 of SML. In May, Axis I was most correlated (in the negative direction) with temperature; Axis 2, with N 2 of the pycnocline, followed by average N2 of SML. In June, Axis I was most correlated (in the positive direction) with temperature; Axis 2, with N 2 of the pycnocline. Chl increased from Group I to Group III.
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Table 1 Average values of physical and biological variables (and range for total Chl) in the SML for the groups of stations (I–III) used in discriminant analyses Variable
April I 3
Total Chl (mg m )
May II
III
I
June II
III
I
II
III
0.18 0.79 2.14 0.67 5.11 14.03 0.61 3.32 11.32 (0.12–0.24) (0.25–1.2) (1.2–3.9) (0.15–0.94) (1.3–11) (11.5–16.8) (0.4–0.8) (1.3–6.2) (9.4–13.7)
MLD (m) 68 N 2 of pycnocline ( 104 s2) 0.61
33 1.33
35 0.55
47 0.88
37 0.66
17 1.46
9 3.73
14 2.91
12 1.69
0.04
0.10
0.05
0.07
0.07
0.17
0.85
0.38
0.57
Salinity
32.94
33.12
33.48
32.92
33.15
33.32
32.73
32.43
32.59
Temperature (1C)
1.77
1.77
1.76
1.75
1.43
0.88
0.92
0.05
0.18
N 2 of SML ( 104 s2)
Table 2 Pooled within-group correlations of physical variables with the two discriminant functions, Axis I and II Variable
MLD N 2 of pycnocline N 2 of SML Salinity Temperature Eigenvalue Cumulative proportion of variance explained
April
May
June
Axis 1
Axis 2
Axis 1
Axis 2
Axis 1
Axis 2
0.24 0.02 0.09 0.58 0.10
0.46 0.97 0.85 0.62 0.11
0.27 0.10 0.22 0.37 0.77
0.19 0.74 0.64 0.45 0.38
0.17 0.28 0.14 0.18 0.34
0.42 –0.72 0.31 0.56 0.17
7.90 0.95
0.43 1.00
5.19 0.98
0.12 1.00
2.47 0.92
0.22 1.00
of the data scatter, respectively): zeu ¼ 16 þ 50e0:531 Chl ðr2 ¼ 0:82; n ¼ 48Þ;
ð6Þ
zeu ¼ 25 þ 40e0:838 Chl ðr2 ¼ 0:78; n ¼ 48Þ:
ð7Þ
At stations close to the lower bound (6), Chl was high near the surface and decreased with depth; at stations close to the upper bound (7), Chl was quite uniform over the euphotic zone (profiles not shown). 3.6. Chlorophyll a and DIN The regression parameters of Chl on DIN changed with location and time (Table 3). After examining scatter plots, the stations were
divided in two clusters (I and II) in April and June, but not in May. Note that the clusters defined here differ from the groupings used for MDA. In April, stations in Cluster I were located on the eastern side (all relations significant, relatively high Chl), while stations in Cluster II were on the western side. In May, the single regression was significant. In June, stations in Cluster I were located along the northern transects (north of 77.351N, plus stations 35 and 60), while stations in Cluster II were along the southern transects (south of 77.01N). The absolute values of the slopes and intercepts increased significantly from April (Cluster I) to May and decreased from May to June (Cluster I) but not significantly.
Z.-P. Mei et al. / Deep-Sea Research II 49 (2002) 4959–4982
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Fig. 8. (a) Average Chl+phaeopigments in the euphotic zone vs. proportion of light absorbed by phytoplankton (equation fitted with the model of Webb et al., 1974: A ¼ 62{1—exp[39 (Chl+phaeopigments)/62]}, r2 ¼ 0:73); and (b) average Chl in the euphotic zone vs. the depth of euphotic zone (negative exponential model), where the two equations are (1) zeu=16+50e0.531 Chl (solid line, r2 ¼ 0:82) and (2) zeu=25+40e0.838 Chl (dashed line, r2 ¼ 0:78).
4. Discussion 4.1. Upper limit of phytoplankton biomass Maximum phytoplankton biomass is determined by availability of resources, such as new nutrients and irradiance, and structure of the food web, given that biomass reflects the balance between production and exports (grazing, sinking,
and lateral advection; Legendre and Le Fe" vre, 1995; Rivkin et al., 1996). We examine here the constraints imposed by nutrients and underwater irradiance. Peak concentrations of chlorophyll in the North Water (ca. 5–15 mg m3 or 160–300 mg m2 at some stations on the eastern side in late May to early June) were as high as or higher than previously reported maxima for the Arctic Ocean.
a
Significance of the slope was tested by permutation; **Po0:01; *0:01oPo0:05; ns (not significant), P > 0:05:
** ** 0.75 0.36 1.04 (1.45, 0.75) 1.80 (2.62, 1.24) 2, 4, 7, 9, 14, 16, 18, 22, 23, 25, 27, 35, 60 31, 33, 36, 38, 40, 44, 49, 50, 52, 54, 68, 72, 76, 80, 82 June
I II
All May
I II April
11.82 (10.46, 13.72) 8.37 (7.18, 10.10)
0.87 1.24 (1.44, 1.07) 12.90 (11.73, 14.16)
0.77 0.10 0.64 (0.98, 0.42) 0.10 (0.06, 0.15) 9, 18, 25, 27, 38, 40, 49, 54 2, 4, 14, 16, 22, 23, 25, 31, 33, 35, 44, 47, 97,
7.40 (5.15, 10.82) 0.86 (1.42, 0.50)
R2 a
b
** ns
Probability of H0 : slope=0a Regression equations Stations Cluster Month
Table 3 Model II linear regressions of average Chl on DIN (NO 3 +NO2 ) in the euphotic zone. In regression equations, boldface values are the intercept (a) and slope (b); values in parentheses are confidence intervals of the coefficients (95% level)
**
Z.-P. Mei et al. / Deep-Sea Research II 49 (2002) 4959–4982
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For example, along a transect from Chukchi Sea to Fram Strait in the Arctic Ocean during July and August of 1994, Chl ranged from 162–445 mg m2 in the Chukchi Sea, comparable to the North Water, to 1–27 mg m2 in the ice-covered Canadian and Amundsen Basins, with values up to 62 mg m2 in the Nansen Basin (Gosselin et al., 1997). In the North Water, the intercept of the regression equation of Chl on DIN (Table 3, May) predicts maximum Chl in the euphotic zone of ca. 13 mg m3 (95% confidence interval of 11.7– 14.2 mg m3), which is lower than the value of 20 mg m3 estimated for the upper 30 m in the same area by Lewis et al. (1996) by assuming that DIN is the limiting nutrient of algal biomass (Lewis et al., 1996; Tremblay et al., 2002). This offset may reflect either interannual variability in initial DIN inventories or the smaller area covered by Lewis et al. (1996; between 761200 and 771200 N over two days in May 1991). The increasing slope of the regression of Chl on DIN from April (Cluster I) to May may indicate increased nitrogen consumption rate by phytoplankton during the bloom peak. Rees et al. (1999) observed that phytoplankton doubled their nitrate uptake rates during the bloom compared with non-bloom phytoplankton. In the North Water, increased grazing activity (i.e. increased ammonium concentration) and sinking of phytoplankton cells (Michel et al., 2002; see also Booth et al., 2002, and Sampei et al., 2002) were likely responsible for loss of phytoplankton biomass in June and July. The use of regressions to estimate the attenuation of light by phytoplankton and non-phytoplankton components (that include water and dissolved and particulate materials) is based on the assumption of constant light attenuation by non-phytoplankton components over the Chl range. This assumption is true for the attenuation due to seawater (0.07 m1; Mitchell and HolmHansen, 1991a and citations therein). In the North Water, the regressions of kd on Chl (Eq. (2)) and on Chl+phaeopigments (Eq. (3)) were very similar, which indicates that the attenuation by phaeopigments was very small. In fact, phaeopigments made up only a minor fraction of the total pigments, especially when Chl was high. This finding indicates that zooplankton grazing was
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relatively low during the phytoplankton bloom, which in turn suggests that non-algal detritus did not significantly contribute to light attenuation in most of the polynya waters sampled. In the North Water, when Chl+phaeopigments reached 5 mg m3, light absorption by phytoplankton in the euphotic zone approached a maximum value of 62% of total light attenuation in the euphotic zone (Fig. 8a), in contrast to oligotrophic systems such as Arctic lakes, with Chl ranging from 0.04 to 1.24 mg m3, where phytoplankton are responsible for 3.2–13.8% of the light absorption and water is the major light-attenuation component (Markager et al., 1999). Similarly, in the Sargarsso Sea, phytoplankton (o1 mg m3 Chl) contribute to 12–39% of the light attenuation in the water column (Smith et al., 1989). According to the data set compiled by Krause-Jensen and Sand-Jensen (1998), when phytoplankton light absorption in the euphotic zone reaches 40%, areal photosynthesis reaches a constant rate because a balance is achieved between further increase in biomass and the consequent reduction in the depth of the euphotic zone. Thus, in the North Water when Chl+phaeopigments were >5 mg m3, the integral rate of phytoplankton production should have been maximum if the depth of the euphotic zone further decreased with increased Chl. The depth of the euphotic zone decreased exponentially with increasing Chl up to 5 mg m3, above which it reached asymptotes of 16 or 25 m (Fig. 8b; discussed further below). As Chl in the euphotic zone increased above 5 mg m3, the fraction of large diatoms (mostly composed of chain- or ribbon-forming diatoms) in >5-mm algae reached 90–100% (Fig. 5 Tremblay et al., 2002). The dominance of large-sized phytoplankton creates a package effect for cellular pigments that leads to lower efficiency of light absorption, thus allowing more light to penetrate deep in the water column (Mitchell and Holm-Hansen, 1991b; Kirk, 1994; Finkel, 1999). The shift in the phytoplankton community to large-sized diatoms could, therefore, explain the constant rather than decreasing depth of the euphotic zone with further Chl increment as predicted by Krause-Jensen and Sand-Jensen (1998; previous paragraph). The
4975
constant depth of the euphotic zone above 5 mg Chl m3 (Fig. 8b) was, to some degree, determined by the vertical distribution of Chl in the water column, which is determined by both phytoplankton growth (depth-dependent) and the strength of vertical mixing. Vertically uniform distributions of phytoplankton in the water column are generally indicative of strong mixing. The stations where Chl values were >5 mg m3 and Chl was uniformly distributed in the euphotic zone fell close to the curve that fits the upper bound of the scatter in Fig. 8b. In contrast, the stations where Chl was concentrated at the surface and decreased with depth in the euphotic zone fell close to the curve that fits the lower bound of the scatter. These patterns indicate that, under similar Chl in a strongly mixed water column, light penetrates deeper than when phytoplankton are concentrated near the surface. Thus, the strength of vertical mixing not only exposes phytoplankton to different time scales of light fluctuations (Denman and Gargett, 1983), but also influences the penetration of light in the water column by modifying the vertical distribution of Chl in the euphotic zone. The maximum biomass of phytoplankton in the euphotic zone of the North Water appears to be determined by the initial nutrient inventory, as suggested by this study and others (Lewis et al., 1996; Tremblay et al., 2002). However, high Chl concentration in the euphotic zone leads to selfshading, which limits population growth well before the maximum biomass is achieved. Therefore, maximum phytoplankton biomass can be achieved when physical conditions (e.g., vertical mixing) and phytoplankton community structure (whether dominated by small or large phytoplankton) allow sufficient penetration of light in the water column, keeping phytoplankton exposed to favorable light conditions. 4.2. Initiation and termination of the phytoplankton bloom The MDA is used here to identify which environmental variables best separate groups of samples with similar Chl concentrations. The three groups of stations for each month were well separated on the first two canonical axes in April
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and May (Fig. 7a and b; the first discriminant function explained 95% and 98% of the total variance, respectively; Table 2). In June, Group I was well separated from Groups II and III, but the separation between the latter two groups was not satisfactory on the second canonical axis. This may be explained by the fact that the lower Chl in Group II than III was caused not only by physical processes but also by export (as observed by Michel et al., 2002). The strong correlation of salinity with the first discriminant function in late April indicates that the higher Chl on the eastern side was associated with high-salinity water (Fig. 7, Tables 1 and 2). The strong correlation of temperature with the first discriminant function in May indicates that the high Chl on the eastern side was associated with relatively high temperature. In June, temperature was the environmental factor mostly correlated with the first discriminant function, which suggests a temperature dependence of Chl consistent with what was observed in May. Indeed, the mean temperature in Groups II and III was 0.051C and 0.181C, respectively, compared with 0.921C in Group I (Table 1). The inverse relationship between stability of the pycnocline at the base of the SML and Chl suggests that the low Chl in Group I stations, which were located on the southernmost transect, was related to the stabilization of the SML (Tables 1, 2 and Fig. 7). By late June, extended stratification and export appears to have prevented a further increase in Chl. It is often stated that a reduction of the MLD leads to increased Chl in the surface water column (e.g., Helbling et al., 1995; Polovina et al., 1995; Arrigo and Weiss, 1998). In the North Water, however, the MLD did not play a major role in discriminating among the Chl-based groups of stations at any time (Table 2). A low correlation of MLD with Chl also was observed in the Southern Ocean (Fig. 5 of Mitchell and Holm-Hansen, 1991a). Several reports have shown that initial stabilization of the surface water column is critical for development of phytoplankton biomass (e.g., Townsend et al., 1992; Strass and Noethig, 1996). Because sampling over the large area of the North Water could not be synoptic, we likely missed the initial phase of stabilization of the SML and the
resulting increase in Chl, at some stations. The low temperature and high correlation between Chl and temperature, however, suggest that the low correlation of MLD with Chl resulted from a limitation of Chl increase by low temperature in April and early May. In May, the strong correlation of temperature with the first discriminant axis suggests that a temperature increase of 0.31C (from Group I to II) to 0.61C (from Group II to III), in the shallow SML, could have enhanced the Chl increase (Tables 1 and 2, Fig. 7). For the Southern Ocean, Tilzer et al. (1986) reported that both phytoplankton photosynthesis and growth increased between 1.51C and 51C. Low temperature in April and early May likely decoupled Chl increase from decreasing MLD. With regards to the debate concerning the ability of the Sverdrup model to predict initiation of the spring phytoplankton bloom, we suggest that the model apply only when temperatures in the SML are above sub-zero. 4.3. Effects of the oceanic heat mechanisms and circulation on phytoplankton dynamics In the North Water, the early spring phytoplankton bloom started on the eastern side, where ice was thinnest. Heat budget estimates indicate that the sensible heat entrained in the SML from depth at stations on the eastern side was not enough to actually melt the ice in April (Melling et al., 2001). Thus, thinner ice on the eastern side could have resulted from slow formation of new ice due to warmer water mass and atmospheric temperatures (Barber et al., 2001). The contribution of ice algae to the initiation of the phytoplankton bloom depends on ice melt (Vincent, 1988; Sakshaug and Skjoldal, 1989). If the disappearance of ice was not caused by ice melt, then the contribution of the release of ice algae to the initiation of the phytoplankton bloom in the North Water could not have been substantial. In any case, the thin ice (Fig. 2) could have favored the early development of phytoplankton in late April and early May, as increased mean irradiance of 30–70 mmol photons m2 s1 was observed in the SML on the eastern side. According to Smith and Harrison (1991), this irradiance level is
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sufficient for shade-adapted phytoplankton to grow in polar regions. Another factor regulating the timing of the phytoplankton bloom may have been lateral advection at the regional scale (Holm-Hansen and Mitchell, 1991; Lucas et al., 1999). High Chl on the eastern side in late April may have resulted from weak northward currents (ca. 5 cm s1; Melling et al., 2001), as Smetacek and Passow (1990) suggest that the critical factor for the spring bloom is the period of stabilization of a shallow layer that must be long enough to permit the accumulation of enough algal cells to overwhelm the populations of grazers. Conversely, a mooring at station N 2 ; located at the northern tip of the polynya on the western side, showed strong southward currents in April 1998 (ca. 10–16 cm s1 at 22 and 102 m depths; Melling et al., 2001). In addition ADCP data from a mooring in the northern part of the polynya (N2, Fig. 1) showed residual velocities of the order of 20–30 cm s1 over the first 100 m in April and May 1998 (Y. Gratton, unpublished data), which may contribute to the slow development of Chl on the western side. Unfortunately, there were too few moorings on the western side of the polynya to examine the possible effects of seasonal variations in lateral advection on Chl. Both the northward flow of Baffin Bay water and the southward flow of Arctic outflow are subject to year-to-year variation, which may have consequences to the timing of the phytoplankton bloom in the North Water. Baffin Bay water flows northward as a weak current (ca.5 cm s1; Melling et al., 2001), which could transport phytoplankton from northern Baffin Bay into the polynya over the months following a bloom there. In Northern Baffin Bay, moderate Chl concentrations (57 mg m2 or 1.26 mg m3), with subsurface chlorophyll maximum of ca. 4 mg m3 and dominant genera of Thalassiosira, Chaetoceros, Rhizosolenia, and Nitzschia, were reported during August and September of 1978 (Harrison et al., 1982). After the productive season, phytoplankton cells may sink and seed the deep Baffin Bay water. Thalassiosira and Chaetoceros also dominated the phytoplankton bloom in the North Water in late April and May 1998 (Booth et al., 2002; Lovejoy et al.,
4977
2002). Given an average speed of ca. 5 cm s1 and the distance to be crossed, the West Greenland Current flows from Baffin Bay to the southeast of the North Water (e.g., stations 54 and 40) in ca. 1– 2 months and to northeast stations (e.g., A18 and A27) in ca. 3–4 months. Hence, if phytoplankton cells did sink to depth in northern Baffin Bay in October, they could have reached the North Water at a time corresponding to the initiation of the phytoplankton bloom on the eastern side of the polynya. Hence, the timing and amount of sensible heat reaching the upper water column from northern Baffin Bay may determine both the extent of ice melt and timing of polynya opening and the possible contribution of ice algae and other taxa to the local phytoplankton community. 4.4. Alternative heat regimes involved in the regulation of phytoplankton biomass The contrasting ice concentrations and the intensity and pattern of currents between the eastern and western sides of the polynya led to the differential timing of the phytoplankton bloom and contrasting patterns of Chl distribution on the two sides of the polynya. High phytoplankton biomass on the eastern side increased light absorption in the water column, and thus decreased the thickness of the euphotic zone (Fig. 8). Since only a small fraction of the incident radiation (usually less than 1%) is used for photosynthesis, the bulk of the absorbed energy is dissipated as heat, which could lead to increased local water temperature and decreased MLD (Platt et al., 1994). Indeed, the heat trapped by phytoplankton has been shown to contribute to the upper-ocean heat budget at both low (Dickey et al., 1998) and high latitudes (Stramska and Dickey, 1993). At high latitudes, experimental and modeling studies have shown that a Chl increase from 0.2 to 2 mg m3 causes a 0.21C increment over 3–7 days (Stramska and Dickey, 1993). In the Arabian Sea, up to 25– 30% of the heating in the upper layer can be attributed to phytoplankton heat absorption when Chl >0.7 mg m3, which enhances the vertical stability of the water column (Dickey et al., 1998). Furthermore, phytoplankton asymmetrically contribute to the upper ocean heat budget by
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diminishing the rate of cooling during the cooling season, but enhancing the rate of heating during the warming season (Lewis et al., 1983; Sathyendranath et al., 1991). In the North Water, direct determination of biological heating effects is not available, because it would have required timecourse monitoring of irradiance, Chl and water temperature at a single station. However, Chl ranging from 1.2 to 3.8 mg m3 at some depths was observed at stations on the eastern side, where daily averaged surface PAR reached B100 mmol photons m2 s1 (after correcting for ice cover), as early as April. In May and June, when the daily averaged surface PAR was >300 mmol photon m2 s1, average Chl in the euphotic zone exceeded 10 mg m3 at stations on the eastern side, which is much higher than in most oceanic environments including those mentioned above (Stramska and Dickey, 1993; Dickey et al., 1998; Krause-Jensen and Sand-Jensen, 1998). Because of the strong increase in solar radiation and Chl in the water column as the season progressed, the trapping of heat by phytoplankton in the upper water column of the North Water should be significant. In September of 1999, Kashino et al. (2002) observed high activity of the diadinoxanthin cycle in the photosynthetic system of phytoplankton, which serves to dissipate excessive light energy as non-radiative heat. All of the available information thus suggest that the early development of phytoplankton could have contributed significantly to the heat budget of the SML in the North Water at a time when air temperature was still too low to warm the upper water column. Such a condition could have contributed to providing warmer temperature for phytoplankton growth and to stabilizing the water column, because part of the heat trapped by phytoplankton also involved heat loss to the atmosphere, preventing water temperature from otherwise decreasing further, even though density is governed predominantly by salinity at low temperature. Thus, an interactive feedback could be expected between increased temperature and Chl, i.e. increased irradiance and temperature initiated phytoplankton growth, after which the increased phytoplankton biomass contributed to the heating of the SML, facilitating the phyto-
plankton bloom on the eastern side from April through June. The above discussion leads to a scenario that explains the early phytoplankton bloom on the eastern side of the North Water. Sensible heat entrained from depth into the SML and warmer air temperature above the region slow new-ice formation, resulting in less ice on the eastern side. The thin ice cover in turn allows higher irradiance in the SML. Along with the seasonal increase in day length and irradiance, increased light attenuation in the shallow euphotic zone, a result of increased Chl, and the heat trapped by the high phytoplankton biomass could facilitate warming and water-column stabilization, favoring the phytoplankton bloom. In contrast, the western side of the polynya was characterized by low Chl, temperature and PAR, and a deep SML in April and May (Table 2). There, deep mixing and low PAR were not conducive to heat trapping in the water column and might delay warming in the predominantly Arctic origin water (Tremblay et al., 2002). In addition, stations on the western side had a strong pycnocline at the base of the SML, with little possibility of sensible heat from depth to be entrained into the SML. The possible input of sensible heat entrained into the SML from deep water and heat trapped by phytoplankton could be critical in accounting for the earlier spring bloom and later accumulation of Chl on the eastern compared to the western side, irrespective of whether sensible heat contributes to the opening of the polynya. The hypothesis supported by this study differs from that of Lewis et al. (1996), who suggested that sensible heat was partly responsible for the polynya opening, the subsequent stratification due to ice melt and the early phytoplankton bloom. Ice melt due to sensible heat input and consequent stratification of the upper water column, however, cannot be verified based on available data.
5. Conclusion The abundance of nutrients contributed to the high phytoplankton biomass in the North Water
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(Tremblay et al., 2002). Without essential physical characteristics, however, such as stabilization of the water column and sufficient light and temperature in the SML, the region could conceivably evolve into a high-nutrient low-Chl (HNLC) system, as described by Dugdale and Wilkerson (1991). The specific physical conditions of the region (Melling et al., 2001; Tremblay et al., 2002), allowing not only sufficient nutrient supply but also temperature and irradiance on different temporal and spatial scales, resulted in the high productivity of phytoplankton during the spring– summer season of 1998. Our study showed that initiation of the phytoplankton bloom in the North Water generally followed the predictions of the Sverdrup model, except that lateral advection and low temperature, especially when lower than 01C, tended to hold the phytoplankton bloom in check. Below 01C, the increase in Chl was decoupled from increased irradiance and reduced MLD. The timing of the phytoplankton bloom has significant implications to the coupling of autotrophic and heterotrophic activities and hence, to the fate of primary production in marine ecosystems (Legendre, 1990; Kirboe, 1993). The differing physical conditions within the North Water not only determined the initiation and development of the phytoplankton bloom, but also created a positive feedback between more heat and increased phytoplankton stock by way of the biological heating effect that contributed to watercolumn stability. In addition, because self-shading started when Chl reached 5 mg m3 in the euphotic zone, well before maximum biomass was achieved, the timing of the phytoplankton maximum was delayed by high phytoplankton biomass later in the season. We have shown here that the spring phytoplankton bloom in the North Water was sensitive to temperatures above 01C, irradiance and lateral advection. Because variations in wind field and atmospheric temperature can modify the timing and extent of ice formation, advection and melting, and thus the strength of water-column mixing and water temperature, the implication is that interannual and long-term changes in physical factors due to climate change will modify the
4979
timing, duration and spatial pattern of the phytoplankton bloom.
Acknowledgements We thank W. Vincent and L. Fortier (Universite! Laval, Canada), and D. Barber (University of Manitoba, Canada) for valuable comments and suggestions. P. Minnett (University of Miami, USA) provided access to the on-deck PAR data set. P. Legendre (University of Montreal, Canada) provided access to his computer program for Model II linear regression analyses. We appreciated the assistance and encouragement from colleagues involved in the overall project and the officers and crew of the Canadian Coast Guard Icebreaker Pierre Radisson during sampling and data acquisition in the field. Special thanks are due to M. Robert for her efficient work at CTD profiling during the whole cruise and to M.-J. Martineau for microscopic examination of the phytoplankton samples. This work was funded by the Natural Sciences and Engineering Research Council of Canada (to L.L., Y.G., M.G., and P.L) and received logistical support from the Polar Continental Shelf Project (Energy, Mines and Resources Canada) and the Canadian Coast Guard. Valuable comments and suggestions from two anonymous reviewers were critical to improving our earlier manuscript. This is a contribution to the research program of GIROQ (Groupe interuniversitaire de recherches oce! anographiques du Que! bec) and to the International North Water Polynya Project (NOW).
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