The Marian Cove is 4.5 km long, 1.5 km wide, ~110 m deep, and is located in the Maxwell Bay, southwest of the King George Island, WAP. The area has experienced ~1.7 km of glacial retreat for 50 years (Park et al., 1998; Lee et al., 2008; Rückamp et al., 2011). Recent (1956–1957 and 2020–2021) glacier retreat velocity was ~27.2 m/yr in the Marian Cove ( Figure 1 ). The phytoplankton biomass (chl-a) in the area started to increase from October to November peaking (i.e., a “bloom”) around January, whereas the lowest values occurred during winter (June–August) (Kang et al., 2002; Lee et al., 2015; Jeon et al., 2021). Marian Cove is characterized by intense summer blooms with high primary production in January (Kim et al., 2021). The general pattern of seasonal phytoplankton succession is that diatoms (>20 µm) dominate in summer and, pico- and nanophytoplanktons (<20 µm) dominate in winter (Kang et al., 1997; Kang et al., 2002; Lee et al., 2015; Jeon et al., 2021).

The study site was conducted from a fixed-point coastal monitoring site for marine ecosystem near King Sejong Station (62°13’ S, 58°47’ W; KSS), Korea ( Figure 1 ). Experiments for C and N uptake by phytoplankton were conducted from December 19, 2018, to January 26, 2019, and results were compared with the environmental parameters (water temperature, salinity, chl-a, nutrients, and phytoplankton taxonomy). The water depth at the site is between 5 and 10 m, and is vertically well-mixed owing to wind- and tide-induced currents (Lee et al., 2015 and references therein). Surface water samples were collected every third day around solar noon at a depth of 0.5 m using a sampler. Average three–hour meteorological data (wind speed and wind direction) were collected during the sampling period from the automatic meteorological observation system (AMOS-3) of the KSS. The wind vane was located 10 m above ground (Park et al., 2013). Ten-minute interval data were averaged into hourly and daily data.

Water temperature and salinity were measured using a YSI Model 30 (Yellow Springs Inc. Ohio, USA). For dissolved inorganic nutrient (NO₂ ^- + NO₃ ^-, NH₄ ⁺, PO₄ ^3-, and SiO₂) concentrations, the seawater was filtered through GF/F filters (pre-combusted at 450°C for 4 h, nominal 0.7 µm, Whatman, UK). The filtered seawater samples were preserved at approximately -20°C until analysis. Major inorganic nutrients were measured using a 4-channel continuous autoanalyzer (QuAAtro, SEAL Analytical, UK), according to the manufacturer’s instructions. Standard curves were run for each sample batch using freshly prepared standards with concentrations in the range of that of the samples. For total chl-a analysis, seawater was filtered through GF/F filters (pre-combusted at 450°C for 4 h). For size-fractionated chl-a analysis, water was passed sequentially through a nucleopore membrane filter (20 µm and 2 µm) and GF/F filters. All filters were extracted overnight with 90% acetone, and the extracts were analyzed using a fluorometer (Trilogy, Turner Designs, Sunnyvale, CA, USA) (Parsons et al., 1984).

The sampled seawater was prefiltered through a 200 µm mesh to remove zooplankton. ¹³C-labeled sodium bicarbonate (Cambridge Isotope Laboratories, USA), ¹⁵N-labeled potassium nitrate and sodium ammonium (Sigma–Aldrich, USA) were added to the surface water samples in 1 L polycarbonate bottles. The ¹³C and ¹⁵N levels were ~4–20% of the total dissolved inorganic C and ambient nitrogenous nutrient concentrations. The bottles were incubated for 4–5 h in ambient seawater. The incubated seawater was filtered (0.3 L) through 450°C pre-combusted GF/F filters (25 mm). A fraction of the picophytoplankton and water passed through 2 μm nucleopore filters (47 mm) and the filtrate was passed through GF/F filters (25 mm). The filters were immediately stored at -80°C until further analysis. The C and N (nitrate and ammonium) uptake rates of the phytoplankton were measured according to the protocol described by Kim et al. (2021). The samples were placed overnight in an acid fume to remove carbonates. C and N isotope abundances were determined using Elemental Analysis (Euro EA3028, EuroVector, Milan, Italy) - Isotope Ratio Mass Spectrometry (Isoprime 100, EIementar, Manchester, UK) in a stable isotope laboratory at Hanyang University, Korea. Carbonate alkalinity of the water sample was determined by titration with 0.01 N HCl, and the total CO₂ content was calculated according to the method of Parson et al. (1984). The C or N uptake rates (mg C (or N) m^-3 h^-1) of phytoplankton were calculated based on Hama et al. (1983) and Dugdale and Goering (1967), respectively, using the following formula (1):

where ΔPOC(orPON)(t) is the increase in particulate organic C or N concentration during incubation (mg C (or N) m^-3), respectively, t is the incubation time, a_is is the atomic % of ¹³C (or ¹⁵N) in the incubated sample, a_ns , is the atomic % of ¹³C (or ¹⁵N) in the natural sample, a_ic is the atomic % of ¹³C (or ¹⁵N) in the total inorganic C (or N). The specific uptake rate can be defined as the rate of uptake or transport of the product (h^-1). Unlike the C uptake samples (duplicates), single samples were analyzed for the N uptake rates.

Pearson’s correlation was used to investigate the relationship between environmental variables and C and N uptake rates. In this study, we estimated the time series of temporally cumulative zonal wind stress (τ_x, in N m^-2) from December 19, 2018 to January 26, 2019, to identify cumulative force exerted on the surface. The cumulative zonal wind stress at a particular time is calculated as the sum of the wind stresses over the previous days (1-day to 5-day in this study). Positive and negative values indicate eastward and westward wind stress, respectively.

The surface water temperature at the sampling site in the Marian Cove was the highest (2.2°C) and lowest (0.3°C) on December 19, 2018 and January 26, 2019, respectively ( Figure 2A ). With the exception of a few days (≤0.6°C) in December to January, most recorded temperatures were above >1°C. The surface salinity ranged from 29.6 to 34.5. Salinity was stable throughout the study period, except a low salinity period in January 12–20. Salinity was positively correlated with the water temperature (r = 0.615, p< 0.05; Table 1 ).

Figure 2B shows the temporal variability of the total chl-a content and the relative contribution of size-fractionated phytoplankton (micro-, nano-, and picophytoplankton). The highest total phytoplankton biomass (chl-a) was recorded on January 6 and 9, 2019 (14.6 and 20.0 mg m^-3, respectively). Average chl-a was 4.4 mg m^-3 (SD = ± 6.0 mg m^-3) during the period from December 19, 2018 to January 26, 2019 ( Figure 2B ). Microphytoplankton constituted the dominant size fraction during the sampling period, accounting for 16.5-93.6% of the total chl-a content, with a mean of 57.2% (± 26.2%), followed by nano- (30.3%), and picophytoplankton (12.5%). At peak chl-a content, microphytoplankton (>20 μm) contributed >90% of the total chl-a. In comparison, the contribution of picophytoplankton (<2 μm) to the total biomass was always low, reaching peak values (28%) on December 25, 2018. The total chl-a concentration was positively correlated with microphytoplankton (p< 0.05, r = 0.748) and negatively correlated with picophytoplankton (p< 0.05, r = -0.651) ( Table 1 ).The daily concentrations of NO₂ ^-+NO₃ ^- and SiO₂ were 11.4–26.2 μM and 41.6–73.5 μM, respectively, although they differed widely on each sampling date. PO₄ ^3- displayed the same pattern as NO₂ ^-+NO₃ ^- ( Figure 2C ). However, the PO₄ ^3- concentration was relatively lower (0.8-1.8 μM). Decrease in nutrient concentrations during sampling were common with substantially increased chl-a concentration, such as from early to mid-January. The largest decrease in nutrient concentration in the surface water from the initial values was observed during January 12–20. NO₂ ^-+NO₃ ^- and PO₄ ^3- concentrations decreased from 19.2 to 11.4 μM and 1.7 to 0.8 μM, respectively ( Figure 2C ). Similarly, low NH₄ ⁺ was observed (<1 μM) at peak chl-a. Interestingly, SiO₂ concentrations further decreased between January 9 and 23, which was not observed for other inorganic nutrients ( Figure 2D ). Overall, the NO₂ ^-+NO₃ ^-+NH₄ ⁺:PO₄ ^3- (N:P) and NO₂ ^-+NO₃ ^-+NH₄ ⁺:SiO₂ (N:Si) molar ratios ranged from 11.7 to 15.4 (average ± SD = 13.2 ± 1.1) and 0.2 to 0.7 (average ± SD = 0.3 ± 0.1), respectively ( Figure 2E ).

Figure 3 shows the three-hourly average wind speed and direction. The average wind speed during the study period was 6 m s^-1. A large change was observed in the average wind speed in the Marian Cove during the sampling period, reaching 14.4 m s^-1. The overall wind direction was WNW during the sampling period. Meanwhile, an easterly wind prevailed from January 10 and 13, with highest wind speeds (>6 m s^-1) ( Figures 3A, B ). The colored curve in Figure 3C shows the accumulated τ_x values, there were two strong eastward episodes. The pulses of accumulated τ_x match well with the wind speed through December 19, 2018 to January 26, 2019 ( Figures 3A, C ). It exhibits clearly that the overall intensity of the eastward in 2-day is the strongest from 1-day to 5-day (see red curve in Figure 3C ). The accumulated intensity of easterly wind in 2-day between January 11 and 13 is the most prominent.

The total C and N (sum of NO₃ ^- and NH₄ ⁺ uptake) uptake rates (mg C (or N) m^-3 h^-1) by phytoplankton in the surface water were measured during summer (December–January) ( Figures 4 , 5 ), and were 0.3–24.5 mg C m^-3 h^-1 and 0.06–3.27 mg N m^-3 h^-1, respectively. The average C and N uptake rates were 6.8 mg C m^-3 h^-1 (SD = ± 7.9 mg C m^-3 h^-1) and 1.0 mg N m^-3 h^-1 (SD = ± 1.1 mg N m^-3 h^-1), respectively, with the highest rates observed on January 9 and the lowest on January 26 ( Figures 4A , 5A ). The temporal pattern of N uptake was very similar to that of C uptake and chl-a content ( Figures 4A , 5A ). Moreover, the specific uptakes (h^-1) of total C and NO₃ ^- were the highest on January 6, after which they gradually decreased, peaked on January 17, and then steadily decreased until the end of the observation period ( Figures 4B , 5C ). Relative contributions of the different N compounds to the measured total N uptake varied for on each sampling date ( Figure 5D ). Compared to total N, proportion of the NO₃ ^- uptake fraction was 15.2% at the beginning (in the middle of December) but increased gradually to 64.7% on January 3.

Although both NO₃ ^- and NH₄ ⁺ were taken up by phytoplankton, the C and dominant N sources varied between the growth phases and blooms. During the sampling period, the 31-fold change in biomass was lower than the 84-fold change in total C uptake, and growth acceleration was much faster in this region. Simultaneously, a more substantial (462-fold) change in the total NO₃ ^- uptake rate was observed, whereas the total NH₄ ⁺ uptake (31-fold) was similar to that of the chl-a content. The total and specific C uptake rates were inversely correlated with DIN (NO₂ ^-+NO₃ ^-+NH₄ ⁺, p< 0.05, r = -0.595, and r = -0.601, respectively) but positively correlated with phytoplankton biomass (chl-a, p< 0.05, r = 0.973, and r = 0.859, respectively) ( Table 1 ). This suggests that the nutrients were alomost simultaneously used by the Marian Cove phytoplankton.

Decrease in the C uptake rate was mainly due to the decline in chl-a values for the total phytoplankton; however, the picophytoplankton fraction showed little variation in chl-a. We found that picophytoplankton productivity remained relatively constant during the sampling period compared to the total uptake, even with nutrient fluctuations ( Figure 4A ). The average surface C uptake by picophytoplankton was 0.2 mg C m^-3 h^-1 (SD = ± 0.2 mg C m^-3 h^-1), the highest being on December 22 (0.7 mg C m^-3 h^-1), and the lowest (0.02 mg C m^-3 h^-1) at the end of the observation period. The contribution of picophytoplankton to the total phytoplankton C uptake ranged from a minimum of 0.6% to a maximum of 27.9% on January 9 ( Figure 4C ). However, the average C uptake by picophytoplankton was up to 8.6% (SD = ± 8.7%) of the total measured C uptake. The specific uptake rate (h^-1) of picophytoplankton showed a similar trend to the uptake rate of C, and decreased with decreasing surface C uptake ( Figure 4B ). A positive relationship was observed between the picophytoplankton C uptake, temperature, and salinity (p< 0.05), while a negative relationship was found between the picophytoplankton relative contribution to total biomass an the total C uptake ( Table 1 ).

NO₃ ^- uptake by picophytoplankton in the surface waters varied from<0.01 to up to 0.04 mg N m^-3 h^-1, with a mean of 0.01 mg N m^-3 h^-1 (SD = ± 0.01 mg N m^-3 h^-1) ( Figure 5E ). The surface NH₄ ⁺ uptake rates varied from 0.005 to 0.257 mg N m^-3 h^-1, whereas the NO₃ ^- uptake rates showed a relatively narrow range (0.001-0.035 mg N m^-3 h^-1) ( Figure 5E ). The NH₄ ⁺ uptake rate was higher than the NO₃ ^- uptake rate in picophytoplankton, suggesting that it preferred NH₄ ⁺ over NO₃ ^- (average NH₄ ⁺ to NO₃ ^- uptake ratio of 9.1). Unlike the C uptake, neither NO₃ ^- nor NH₄ ⁺ uptake showed any trend (data not shown). No statistical relationship was observed between picophytoplankton N productivity and hydrography and/or dissolved inorganic nutrient concentrations ( Table 1 ).

Relatively high phytoplankton biomass (up to 45 mg m^-3 near-shore waters adjacent to Palmer station; Goldman et al., 2014) and productivity have been observed in coastal and shelf areas (Garibotti et al., 2003; Arrigo et al., 2008; Ducklow et al., 2012b; Mendes et al., 2012). Particularly during the summer growing period, the phytoplankton species composition varied seasonally, with species abundance related to changes in environmental conditions based on 15 years of monitoring data (1996–2011) (Jeon et al., 2021). For example, seasonal nutrient depletion reflects diatom-dominated phytoplankton communities in regions where blooms are common, with a significant contribution of diatoms to the total phytoplankton biomass in the WAP (Clarke et al., 2008; Vernet et al., 2008; reviewed in Ducklow et al., 2012a; Höfer et al., 2019, and references therein).

Similarly, we observed massive phytoplankton blooms (chl-a up to 20 mg m^-3), which were mainly composed of microphytoplankton (>20 µm) and large diatoms (>40 µm, e.g., Thalassiosira spp.; Supplementary Figure 1 ). There was a strong relationship between microphytoplankton fraction and the total chl-a concentration (p< 0.05, r = 0.748, n = 13), where microphytoplankton contributed >90% of the total chl-a when the latter exceeded ~14 mg m^-3. When the chl-a concentration was<2 mg m^-3, nanophytoplankton was occasionally the dominant size class ( Figure 2B ). This is consistent with Jeon et al. (2021), who reported that centric and pennate diatoms comprised a substantial part of phytoplankton biomass in the Marian Cove, with an ~58% contribution to the total microphytoplankton abundance during summer blooms. Therefore, seasonal blooms result primarily from increases in microphytoplankton and diatoms.We observed similar trends in the C uptake data and chl-a concentrations. During the peak bloom, the C uptake rate was much higher (24.5 mg C m^-3 h^-1) and the macronutrient concentrations declined considerably, implying a much more active nutrient intake by the phytoplankton. Variability in macronutrient consumption is reflected in the phytoplankton biomass in the Antarctic coastal waters during summer (Clarke et al., 2008; Zhang et al., 2019, and references therein). The dominance of microphytoplankton is consistent with the total primary production, and also positively correlates with the chl-a in various coastal waters (Varela et al., 2002; Marañón et al., 2012; Kim et al., 2021). Previous studies have documented a markedly high chl-a and C uptake rate in the coastal ice-edge zone in Antarctic waters adjoining the Indian Ocean (up to 4.0 mg m^-3 and 3.3 mg C m^-3 h^-1, respectively; Verlencar et al., 1990), Marguerite Bay (>10 mg m^-3 and >4.2 mg C m^-3 h^-1, respectively; Ducklow et al., 2012b), Marian Cove (up to 19.5 mg m^-3 and 31.1 mg C m^-3 h^-1, respectively; Kim et al., 2021), and South Bay (>15 mg m^-3 and >50 mg C m^-3 h^-1, respectively; Höfer et al., 2019). This is consistent with observations from the Antarctic coastal waters, and supports the theory that this is one of the most important global “biological hotspots”. The development of phytoplankton blooms is also an important factor affecting primary production.However, the observed temporal variability (0.3–24.5 mg C m^-3 h^-1) probably reflects a combination of changes in light availability due to turbidity and water stability (Schloss et al., 2002; Kim et al., 2021), as well as water movement (Höfer et al., 2019). Although various environmental factors affected the C uptake rate, no statistically significant differences were observed in the direct light intensity, SPM, water temperature, and salinity during the sampling period. Low macronutrient contents were observed during the peak bloom, but the N/P ratio remained similar regardless of the growth phase and was unlikely to act as a limiting factor for phytoplankton growth. Earlier studies reported that macronutrient values in the Marian Cove are generally a non-limiting factor for phytoplankton growth (Lee et al., 2015; Jeon et al., 2021), both due to inflow from the Maxwell Bay and the penguin colonies (Pruszak, 1980; N&eogon;dzarek, 2008), where turbulent mixing and wind-driven upwelling cause continuous and abundant nutrient flow into the photic layer. Additionally, lateral iron inputs from glacial meltwaters and terrestrial sources are sufficient to sustain large-scale phytoplankton blooms on the King George Island (Schloss et al., 2014; Kim et al., 2015; Höfer et al., 2019), and phytoplankton growth is not limited, even during intense blooms (Bown et al., 2017).

Nevertheless, the significant reduction in biomass and total C uptake immediately after peak chl-a (on January 12) may have enhanced other phytoplankton losses, such as zooplankton predation and other loss rates (i.e., viral lysis, sinking, and aggregation) (Petrou et al., 2016). Grazing effects, phytoplankton sinking, and mortality rates were not explored in this study. Instead, heterotrophic activity and the physiological status of phytoplankton were estimated based on the NH₄ ⁺ concentration and specific C uptake by phytoplankton. For example, being its main contributor, zooplankton excretion can boost ambient NH₄ ⁺ concentrations (Corner and Davies, 1971). Bacterial degradation of organic matter may also result in summertime NH₄ ⁺ accumulation (Koike et al., 1986). NH₄ ⁺ and NO₃ ^- concentrations in the area showed a rapid increase under lower C uptake on January 12 ( Figures 3D and 4A ). Simultaneously, the specific C uptake rates by phytoplankton were approximately 1/6 of the peak value (0.03 h^-1). These results imply a physiological limitation of phytoplankton growth on January 12, which may have temporarily enhanced nutrient concentrations in the surface water due to less consumption by phytoplankton. This suggests the increased possibility of the influence of external forces (i.e., bottom-up control) rather than biointeractions (i.e., top-down control).

Utilisation of inorganic N by phytoplankton was observed to be highly variable in the Marian Cove during summer. The total NO₃ ^- uptake was higher than the NH₄ ⁺ uptake in the bloom phase, while primary production in the system evolved from NO₃ ^–-based to NH₄ ⁺-based at the end of sampling. This evolution is in agreement with earlier studies that reported rise in NH₄ ⁺ levels during summers, when zooplankton and microbial metabolism provide regenerated N (Clarke et al., 2008; Mosseri et al., 2008). The high ambient NH₄ ⁺ concentrations in conjunction with increased NH₄ ⁺ uptake are indicative of an active regeneration process in the Marian Cove and are attributed to the late summer sampling. Thus, we hypothesized that NH₄ ⁺ uptake occurs in relation to changes in its concentration and phytoplankton species/size composition.

Dependence of the NH₄ ⁺ concentration on NO₃ ^- uptake has been described in previous studies (Dortch, 1990; L’Helguen et al., 2008; Dugdale et al., 2012; Glibert et al., 2016). For example, NH₄ ⁺ concentrations >1 µM (Dortch, 1990) or over a certain threshold (0.2–100 µM; Varela and Harrison, 1999; Dugdale et al., 2012; Glibert et al., 2016) inhibit NO₃ ^- uptake rates. Our data also support that a relatively higher uptake of NH₄ ⁺ than that of NO₃ ^- was observed in the presence of >1 µM NH₄ ⁺ in ambient waters, which might be influenced by the inhibition of assimilated NO₃ ^- by phytoplankton. However, the elevated NH₄ ⁺ uptake did not enhance the total phytoplankton C uptake or biomass. The observed rapid NO₂ ^-+NO₃ ^- drawdown and uptake rate of NO₃ ^- in January, coupled with increases in chl-a and C uptake, indicate high productivity. These results clearly suggest that NO₃ ^- is the preferred growth- and biomass-promoting N source for the studied phytoplankton community. Notably, the NH₄ ⁺ uptake by phytoplankton was higher despite the high NO₃ ^- concentrations. Consequently, inhibition of NO₃ ^- uptake by NH₄ ⁺ (>1 µM) may be one of the factors that led to the low f-ratio (0.08–0.24, Figure 5B ) in the total phytoplankton (f-ratio= NO₃ ^- uptake/total N uptake (NO₃ ^- + NH₄ ⁺ uptake), Eppley and Peterson, 1979).Another possible cause of the lower f-ratio could be the increasing contribution of the smaller cells (<20 µm). Microphytoplankton levels were higher during relatively high biomass and C and NO₃ ^- uptake rates, reaching an f-ratio of up to 0.67 on January 9 ( Figure 5B ). The NH₄ ⁺ uptake rates were the highest in the picophytoplankton fraction throughout the sampling period, resulting in an f-ratio<0.5, which decreased to 0.03 at the end of the sampling period ( Figure 5E ). In picophytoplankton, NH₄ ⁺ is preferentially assimilated over NO₃ ^- because NO₃ ^- assimilation is more energy consuming than NH₄ ⁺ assimilation. Once NO₃ ^- is transported into the cell, it must be further reduced to NH₄ ⁺ before it can be assimilated (Mulholland and Lomas, 2008).More importantly, f-ratio values show that planktonic diatoms (larger cells) are replaced by small phytoplankton in surface water, causing decreased primary production and C export. This is because NO₃ ^–-fueled new production at steady state is equivalent to the organic C that can be exported from total production in the euphotic layer (Eppley and Peterson, 1979). In this regard, phytoplankton species shifts toward smaller values could substantially lower C export efficiency as well as primary production (Montes-Hugo et al., 2008; Montes-Hugo et al., 2009; Lee et al., 2015, Mendes et al., 2012; Rozema et al., 2017; Schofield et al., 2017). The undergoing reductions in diatom silica production in response to ocean acidification and shifts toward smaller cells could reduce the vertical fluxes of diatoms and diminish C export efficiency before the end of this century (Petrou et al., 2019). Therefore, in the neritic area of Marian Cove, nutrient utilization in the surface layer during the summer induced variations in phytoplankton size composition in terms of chl-a and C and N uptake, from summer blooms (early January) skewed toward the microsized fraction (>20 μm) undergoing a gradual shift toward the pico sized fraction (<2 μm) through the nanosized fraction (2–20 μm) in late January. This shift in the size of the phytoplankton community probably involved the prevailing wind direction associated with physiological and morphological properties (e.g., nutrient uptake and export rate), as mentioned above.

Surface wind, meltwater dynamics, and currents, as key drivers of the upper water layer structure, strongly influence nutrient availability and phytoplankton growth (Deppeler and Davidson, 2017). Wind speed and direction may influence the strength and duration of phytoplankton blooms (Schloss et al., 2014; Deppeler and Davidson, 2017, and references therein). High-frequency (HF) wind forcing is important for the residence time of water (Kohut et al., 2018). The major component of HF zonal wind forcing was the so-called easterly wind events, which played a critical role in modulating water movement. Previous, studies on the Marian Cove (Yoo et al., 1999 and Yoo et al., 2015; Jeon et al., 2021) and the adjacent Potter Cove (K1öser et al., 1994) and Admiralty Bay (Nowosielski, 1980) have suggested that the prevailing easterly wind can generate outflow and upwelling near the cove head or deter the flow rate of surface influents from the Maxwell Bay. Pruszak (1980) observed the drift ice movement and reported that, surface water circulation can be controlled with a wind speed of >4 m s^-1 in the adjacent Admiralty Bay. There is also a study result that surface water flows out of the Marian Cove within one day when the east wind blows with an average wind speed of over 8 m s^-1 (Yoo et al., 1999).

Similarly, in our study, the wind direction changed from west to east, and continued for approximately three days. Additionally, strong winds (>6 m s^-1) blew simultaneously, which is believed to have caused low biomass and C and N uptake rates. Although easterly winds occurred on January 6, their speed was low compared to that during January 10–12. In detail, Figure 3C shows the evolution of zonal wind stress during sampling period. As a result of the analysis of cumulative wind stress from Day 1 to Day 5, the wind forcing was higher when 2 days were considered. Overall, negative ∑τ _x values with low wind speeds led to high chl-a, C, and N uptakes from January 2 to 9. This is because factors, such as outflowing water from the Marian Cove to the Maxwell Bay, also play a role in transporting particulate matter, including algae, near the ice wall and resuspending it (Brandini and Rebello, 1994; Yoo et al., 1999). This surface circulation influenced by the easterly wind can deliver surface water and suspended particles to the Maxwell Bay because the Coriolis effect is not important for a small bay width (~1.5 km for the Marian Cove) (Yoo et al., 2015). Resuspension of sediments and benthic diatoms in the inner Admiralty Bay induced by wind-driven upwelling (<20 m water depth) increases surface turbidity and nutrient concentrations (Brandini and Rebello, 1994); however, the effect of resuspension was not significant in the present study. As the SiO₂ concentration was lower than that at peak bloom, benthic diatom species (Cocconeis spp.) accounted for 8.5% of total diatom abundance ( Supplementary Figure 1 ). On January 12, the dominant species were small (<20 µm) Navicula spp. and Fragilariopsis spp., accounting for 72.2% to the total diatom abundance that lives mainly on the ice walls and in low-salinity coastal waters (Kang et al., 1999; Fernandes and Procopiak, 2003; Jeon et al., 2021). More specifically, a change in the structure of the phytoplankton community was observed during the sampling period in microscope-based phytoplankton observation ( Supplementary Figure 1 ). Phytoplankton communities were dominated by bacillariophytes (diatoms), which comprised >60% (up to 96.3%) of the total C biomass over the sampling period. Thalassiosira spp., accounting for 67.7% of total diatom abundance at peak bloom, decreased to 1/4 of its peak bloom value on January 12 ( Supplementary Figure 1 ).Water column stratification and relatively weak wind speed are favorable for initiating seasonal blooms, and the presence of sea ice and glacial meltwaters influences primary production patterns (Saba et al., 2014; Rozema et al., 2017; Höfer et al., 2019; Kim et al., 2021). In this regard, the relatively weak wind speeds observed at the beginning of January seem to enable water column stabilization, favoring an increase in the residence times of cells in the photic layer and providing more light for photosynthesis, allowing high phytoplankton accumulation in the cove.