Iron and Phytoplankton Growth in the Subarctic North Pacific

Shigenobu Takeda

Department of Aquatic Bioscience, Graduate School of Agricultural and Life Sciences, The University of Tokyo, Yayoi 1-1-1, Bunkyo-ku, Tokyo 113-8657, Japan

Abstract

Iron availability has been shown to have potential controls on phytoplankton growth, nutrient utilization, algal community composition, and the ecosystem structure in the subarctic North Pacific. Recent findings on the lateral iron transport from the surrounding marginal regions to the pelagic waters highlighted the importance of particulate iron in the subarctic North Pacific, but the transformation between dissolved and particulate phases and its interaction with organic ligands are still uncertain. In spite of active researches in the subarctic high-nitrate, low-chlorophyll (HNLC) waters, significant impacts of Asian dust on the phytoplankton productivity have not been detected, suggesting spatial and temporal mismatch between the dust inputs and biological activities. Satisfaction of algal demands for both light and iron is a key for phytoplankton blooming in the HNLC waters. Surprisingly, the community half-saturation constant for growth with respect to iron was found to be similar between the western and eastern gyres; however, differences in the iron supply process and its availability in these two gyres seem to have developed unique phytoplankton populations. It is essential to evaluate iron transport processes that work on a time-scale needed for phytoplankton blooms, and further studies are needed at the central regions of the subarctic North Pacific.

Keywords

iron, organic ligands, silicic acid, nitrate, light, temperature, phytoplankton, diatoms, iron fertilization, HNLC, Western Subarctic Gyre, Alaska Gyre, North Pacific


Received on September 7, 2009

Accepted on April 28, 2011

Published online on October 15, 2011

*Present address:

Faculty of Fisheries, Nagasaki University, Bunkyo-machi 1-14, Nagasaki 852-8521, Japan.

e-mail: s-takeda@nagasaki-u.ac.jp


1. Introduction

Phytoplankton growth in the oceans requires many physical, chemical, and biological factors. The major plant nutrients in the ocean, commonly thought to be critical for phytoplankton growth, are nitrate, phosphate, and silicic acid that exist at micromolar levels. A number of trace metals are also essential for the growth of the organisms. These elements function as active centers or structural factors in enzymes and electron carrier proteins. Iron is required in the largest amounts than any of the trace metals. It has been suggested that cells evolved high iron requirements during the early history of life, before the evolution of oxygenic photosynthesis, when iron was probably one of the major ions in seawater. However, solubility of iron in seawater decreased by six or seven orders of magnitude at the anoxic/oxic transition (Brand 1991a). Then, it became problematic for marine primary producers to obtain adequate iron in seawater. In most of the present-day open oceans, dissolved iron is often found at concentrations of only subnanomolar level in surface waters (Landing and Bruland 1987; Johnson et al. 1997), owing to both inorganic precipitation and biological uptake.

The discovery of extremely low iron concentrations in surface waters and its nutrient-like vertical profiles led John Martin, Director of the Moss Landing Marine Laboratory, to speculate that phytoplankton growth is limited by the availability of iron in some areas of the ocean. Both diatoms and small phytoplankton become iron-limited in large oceanic regions during summer, mainly in high-nitrate, low-chlorophyll (HNLC) waters, which are thought to represent about 40–50% of the areal extent of the world's oceans (Moore et al. 2002). The subarctic North Pacific, Eastern equatorial Pacific, and the Southern Ocean are the major HNLC regions, having high macronutrient concentrations, adequate light, and physical characteristics required for phytoplankton growth in the summer but with very low phytoplankton biomass.

By applying careful sampling and analytical techniques to analyze trace metals in seawater, John Martin conducted onboard bottle incubation experiments using surface waters of the eastern subarctic North Pacific, and these experiments demonstrated dramatic phytoplankton growth upon the addition of nanomolar iron, when compared with the control bottles in which no iron was added (Martin and Fitzwater 1988). His results renewed the scientific debate over what controls phytoplankton production in nutrient-rich areas of the open sea (Chisholm and Morel 1991). John Martin's Iron Hypothesis linked iron supply, rates of primary production, and subsequent carbon sequestration in HNLC waters both in the present day and the geological past (Martin 1990). Furthermore, the hypothesis put forward by Morel et al. (1991) offers an explanation for the widespread existence of HNLC waters. Based on low iron supply constraining the growth rate of large phytoplankton and grazer control keeping small phytoplankton cropped to low biomass levels, this hypothesis links iron supply, food-web structure, and macronutrient cycling.

The subarctic North Pacific has two prominent gyres—the Alaska Gyre and the Western Subarctic Gyre (Fig. 1). These two gyres are characterized by relatively uniform distributions of temperature, salinity, macronutrients, and light, yet they have strong zonal gradients in atmospheric iron (dust) deposition transported eastward from the Asian desert and loess regions (Duce and Tindale 1991). Most of the dust reaches the North Pacific during springtime, and therefore, iron could be a limiting factor for phytoplankton growth and primary productivity both in the Alaska Gyre and the Western Subarctic Gyre during the low dust season in summer and fall, but not in the winter when iron and light may be co-limiting (Maldonado et al. 1999). The difference in episodic iron deposition is presumed to give rise to distinct phytoplankton communities, that is, centric diatoms in the Western Subarctic Gyre versus pennate diatoms in the Alaska Gyre, which characterize these biogeochemical provinces (Harrison et al. 1999).


Fig. 1. General circulation in the Subarctic North Pacific showing the Alaska Gyre and Western Subarctic Gyre. Double arrows are intense boundary currents. The Subarctic Boundary separates the subarctic Pacific region to the north from the subtropical Pacific region to the south. Reprinted from Prog. Oceanogr., 43, Harrison et al., Comparison of factors controlling phytoplankton productivity in the NE and NW subarctic Pacific Gyres, 205–234, © 1999, with permission from Elsevier.

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Early iron-enrichment bottle incubation experiments provided evidence for iron limitation both in the Alaska Gyre and the Western Subarctic Gyre (Martin and Fitzwater 1988; Boyd et al. 1996; Takeda 1998). However, criticism that such small-scale, enclosed experiments may not accurately reflect the response of the plankton community has led to three mesoscale in situ iron-enrichment experiments in the subarctic North Pacific (Tsuda et al. 2003, 2007; Boyd et al. 2004). These experiments demonstrated that iron availability strongly influences primary productivity and food-web structure in summer. On the other hand, the observed strong grazing control on phytoplankton responses in one of these experiments emphasized the complexity of ecosystem responses to iron input in the subarctic North Pacific.

It is clear that our understanding of iron biogeochemistry has advanced substantially over the past two decades, and has resulted in a new paradigm that considers iron as one of the nutrients similar to nitrate and phosphate, although iron chemistry in seawater and its biological uptake mechanisms are very complicated. In this monograph, the author begins with what we know about the distribution and fluxes of iron in the subarctic North Pacific, then discusses iron requirement for phytoplankton growth, and examines their responses to iron addition in the bottles under several different conditions in the Northwest (NW) and Northeast (NE) Pacific. Lastly, the key biological and geochemical results obtained in the three mesoscale experiments are summarized, which provide new insights on the roles that iron plays in phytoplankton physiology and ecology in the subarctic North Pacific Ocean.

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2. Distributions of iron and its possible sources in the North Pacific

As iron is one of the important factors controlling phytoplankton productivity in the subarctic North Pacific, there are considerable interests in determining the distribution of iron and its sources. A previous study indicated that the deposition flux of iron-containing dust over the western North Pacific is an order of magnitude higher than that in the eastern regions (Duce and Tindale 1991). On the other hand, recent studies showed that dust fluxes, which were estimated based on dissolved aluminum concentrations in surface water, were lower than those of the previous estimates (Measures et al. 2005), and the supply of dissolved iron from deeper layer to the surface dominated north of 45°N in the subarctic NW Pacific (Brown et al. 2005). In addition, re-suspended iron-containing particles from continental shelf sediments could be transported over long distances by eddies and water current systems (Johnson et al. 2005; Lam et al. 2006). However, relative importance of these iron supply processes on iron distribution and phytoplankton production has not been quantitatively evaluated due to limited information on the spatiotemporal variations of iron in the North Pacific.

In this section, the author overviews the differences in the vertical distributions of iron between the eastern and western subarctic North Pacific, the seasonal variability of dissolved iron concentrations, and the influences of possible iron sources on the distribution of iron in the subarctic North Pacific to provide the hint to extract a key process affecting plankton ecosystems in these waters.

2-1. Sampling and analytical methods of measuring iron in seawater

Details of the sampling and analytical methods of iron determination in seawater have been described in previous papers (Nishioka et al. 2001, 2003, 2007). Basically, seawater samples were collected using acid-cleaned Teflon-coated Niskin-X or Go-Flo bottles attached to a Kevlar wire or a clean CTD-carousel multiple sampler. The filtered samples were obtained by passing through 0.22-μm Durapore filters connected to the spigot of the sampler bottle under gravity pressure. The filtrate and unfiltered samples were adjusted to pH 3.2 and analyzed by automated chelating resin concentration and chemiluminescence detection system (Obata et al. 1993, 1997).

In this study, "dissolved iron" is leachable iron in 0.22-μm filtrate at pH 3.2, and "total dissolvable iron" is dissolved plus leachable iron in unfiltered sample at pH 3.2. The dissolved iron was further size-fractionated by using a 200-kDa polyethylene hollow-fiber ultrafilter unit, and the measured concentrations in the filtrate are defined as "soluble iron." The concentration of the colloidal (200 kDa to 0.22 μm) iron was calculated from the difference in the measured concentrations of "dissolved iron" and "soluble iron." In some studies, unfiltered samples were adjusted to pH < 1.8 and stored for more than 1 year before the analyses, and the measured concentrations were treated as "total iron." The detection limit was 0.017–0.032 nM among the cruises. When analyzed with the reference standard seawater from SAFe cruise (Johnson et al. 2007), this iron measurement method gave a good agreement with the certificated values (Nishioka et al. 2007).

Concentrations of iron(III)-complexing organic ligands were determined by competitive ligand equilibration-adsorptive cathodic stripping voltammetric (CLE-ACSV) method according to Kondo et al. (2007). TAC (2-(2-thiazolylazo)-p-cresol) was used as a competitive ligand for the CLE-ACSV analyses (Croot and Johansson 2000). Ligand concentrations and their conditional stability constants were calculated using Langmuir transformation. The detection range of conditional stability constants for ligands was found to be highly dependent on the competing ligand used and the respective concentrations chosen for the analysis. The value of conditional stability constant for Fe(TAC)2 was observed to be 1022.4, and the center of the detection window in this study was 250. The obtained conditional stability constant was an average value of all the detectable ligands in the sample. The coefficient of variation of ligand concentration measured for 1 nM desferrioxamine B dissolved in UV-irradiated seawater was 5% (n = 3).

2-2. Vertical distributions of iron in the eastern and western subarctic North Pacific

In the eastern subarctic North Pacific, very low dissolved iron concentration of ∼0.05 nM in surface waters and uniform values of 0.62–0.70 nM in the deep-water column, the so-called nutrient-type distribution, had been reported at Ocean Station P (50°N, 145°W; Martin et al. 1989). The apparently close correlation with the nitrate distribution suggests the strong coupling of iron with the oceanic biological cycle. At Ocean Station P in September 1998, there was extremely low iron in colloidal (200 kDa-0.22 μm) and labile particulate (>0.22 μm) fractions in the surface mixed layer (Fig. 2). The soluble iron (<200 kDa) concentrations were 0.06–0.07 nM in the surface mixed layer and represented >85% of dissolved iron above the 40-m depth. Higher concentrations of soluble iron (0.18–0.44 nM) and colloidal iron (0.14–0.20 nM) were observed at 200–800-m depth. In this deep layer, the colloidal iron represented 24–47% of the dissolved iron. Similarity in the increasing pattern with depth between the colloidal iron and macronutrients suggests that the increase in colloidal iron concentration below the surface mixed layer was caused by the formation of iron-containing colloidal particles by remineralization of biogenic sinking organic particles (Nishioka et al. 2001). On the other hand, labile particulate iron concentrations were very low even in the deep waters. Takata et al. (2006) also found extremely low concentrations of labile particulate iron throughout the water column at 41°–45°30'N along 165°W, but there was high concentration of labile particulate iron at 50°N, 165°W, probably due to the input of iron from the Alaskan continental margin to the Alaskan Stream (Martin et al. 1989).


Fig. 2. Vertical distribution of size-fractionated iron in the western subarctic North Pacific (WSNP) and the eastern subarctic North Pacific (ESNP). (a) The data were collected at Oyashio region (49°N; 157°30'E) in May 2000; observed surface mixed layer was 30 m. (b) The data were collected at St. KNOT (44°N; 155°E) in May 2000; observed surface mixed layer was 10 m. (c) The data were collected at Ocean Station P in September 1998; observed surface mixed layer was 40 m. Nishioka et al., Size-fractionated iron distributions and iron-limitation processes in the subarctic NW Pacific, Geophys. Res. Lett., 30(14), 1730, © 2003 American Geophysical Union. Reproduced by permission of American Geophysical Union.

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In the western subarctic North Pacific, the dissolved iron concentrations in the surface mixed layer were 0.08–0.15 nM during late spring, with minimum concentrations of 0.05–0.11 nM observed at 20–40-m depth (Fig. 2). Below 100 m, the dissolved iron concentrations increased and reached a maximum of 1.41–1.66 nM at 600–800-m depth. Soluble iron concentrations in the surface mixed layer were 0.06–0.15 nM. Below 100-m depth, the soluble and colloidal iron concentrations increased sharply. In contrast, iron concentrations in the labile particle fraction were high from the surface mixed layer to deep waters in the western subarctic North Pacific, and its contribution to the total iron was much higher at Oyashio region, when compared with that at Station KNOT in the Western Subarctic Gyre. Recent studies also found that high concentrations of labile particulate iron in the surface water can be observed only in subarctic cold water masses north of the Subarctic Front, which was defined as 6°C (Nishioka et al. 2007), and that labile particle iron exists at high concentration levels in the deep waters at 41°–47°N along 165°E (Takata et al. 2006). In the central part of the subarctic North Pacific (38°30'–47°30'N; 175°30'E), Kuma et al. (1998) reported dissolved (<0.45 μm) iron concentrations in the range of 0.25–1.04 nM, and their surface concentrations were relatively high, when compared with those in the eastern subarctic region.

Comparison of vertical profiles of size-fractionated iron between the eastern and western subarctic North Pacific clearly showed higher concentrations of labile particulate iron toward the west, and this result strongly supports the higher iron supply in the western region. This trend has been significantly observed not only for these three stations but also for the statistically processed 61 data from 12 stations (P < 0.01) (Table 1). Magnitudes of the increases in both colloidal and soluble iron concentrations with depth from subsurface to intermediate water were also greater at the western regions. Greater increased gradient in dissolved iron with depth from subsurface to intermediate water in the western region, relative to that of eastern region, as well as high particulate iron in the whole water column have also been supported by other observations (Fujishima et al. 2001; Brown et al. 2005; Kinugasa et al. 2005; Takata et al. 2006; Nishioka et al. 2007). Therefore, it is important to recognize the existence of ∼3 times higher concentration of labile particulate iron in the western subarctic seawater for better understanding of the iron dynamics in the North Pacific. Indeed, there is an east-west gradient in the photosynthetic competence of phytoplankton between the Western Subarctic Gyre and Alaska Gyre, probably due to the high iron levels in the western region (Suzuki et al. 2002). On the other hand, dissolved iron concentrations in the surface mixed layer in the western region were found to be as low as those in the eastern region. Nishioka et al. (2003) argued that low dissolved iron in the western subarctic North Pacific can be explained by the conversion of colloidal iron to labile particulate iron by aggregation and adsorption onto the suspended organic particles. However, reactivity of colloidal iron with free iron-complexing organic ligands, which exist in excess of soluble iron, should be examined to determine the fate of colloidal iron in the surface water.


Table 1. Comparison of iron concentration in labile particulate and dissolved fractions in the surface mixed layer between the Western and Eastern Subarctic North Pacific. Nishioka et al., Size-fractionated iron distributions and iron-limitation processes in the subarctic NW Pacific, Geophys. Res. Lett., 30(14), 1730, © 2003 American Geophysical Union. Reproduced by permission of American Geophysical Union.

aThe data are shown as the average value ± the standard deviation calculated for labile particulate and dissolved iron concentrations in the surface mixed layer at stations in Oyashio region (May–June 2000). bIn WSNP region (May–June 2000, and July–Aug. 2001). cIn ESNP (Sep. 1997 and Feb., June, and Sep. 1998).

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While iron has thermodynamically very low solubility at the slightly basic pH of oxic seawater, there is recent evidence showing that the bulk of the dissolved iron in the open ocean is complexed with dissolved organic ligands (Van den Berg 1995; Rue and Bruland 1995). Low concentration of dissolved iron present in surface waters, coupled with the excess of strong organic ligands, results in extremely low equilibrium concentrations of dissolved inorganic iron. In the subarctic NE Pacific (50°N; 160°W), a vertical profile of dissolved iron and iron-complexing organic ligand concentrations suggests that most of the dissolved iron has been complexed with organic ligands in the water column, but the concentrations of dissolved iron around 1000–2000-m depth have been higher than the organic ligand concentrations (Fig. 3). The excess dissolved iron can exist as colloidal iron and/or organic/inorganic complexes with weak ligands that could not be detected by this method.


Fig. 3. (a) Vertical profile of dissolved iron (D-Fe), iron-complexing organic ligand concentrations, and (b) conditional stability constant with respect to inorganic Fe' (K'Fe'L) measured at 50°N, 160°W (Station 13, KH-05-2 cruise) in the eastern subarctic Pacific. The error bars show the analytical errors (Kondo and Takeda, unpublished).

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The relationship between iron-complexing organic ligands and dissolved iron concentrations in the subarctic Pacific is shown in Fig. 4. The low concentration of dissolved iron was observed with an excess of strong organic ligands, and most of the dissolved iron was estimated to be complexed with these organic ligands in the surface water. In the deep waters, in general, iron-complexing organic ligands and dissolved iron concentrations exist at almost 1:1 ratio. The average values of the conditional stability constant with respect to inorganic iron species (log K'Fe'L) were 12.1 ± 0.53 M–1 for 10-m waters (n = 24) and 12.7 ± 0.25 M–1 for deeper waters (n = 6). Similar relationship between Fe(III) hydroxide solubility and labile dissolved iron concentrations had been observed in the western North Pacific (Kuma et al. 2003; Takata et al. 2004). These results suggest that the concentration of iron-complexing organic ligands varies in the surface water and could have a strong influence on the chemical speciation of dissolved iron, while dissolved iron in deep water seems to be under the control of the complexation capacity of the iron-complexing organic ligands.


Fig. 4. Comparison of iron-complexing organic ligands with dissolved iron concentrations in (a) surface water (10 m) and (b) deep water (75–5000 m) measured in the subarctic North Pacific. The dotted line shows 1:1 ratio (Kondo and Takeda, unpublished).

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2-3. Temporal heterogeneity of iron distribution

In the eastern subarctic North Pacific, spatial and temporal changes in vertical distribution of size-fractionated iron were studied during September 1997, June and September 1998, and February 1999 along a transect extending from southern Vancouver Island to Ocean Station P (Line P; Nishioka et al. 2001). At Ocean Station P, vertical profiles of dissolved iron exhibited nutrient-like distributions (Fig. 5). The observed profiles showed good agreement with the previously reported data by Martin et al. (1989), except February 1999. In February 1999, the dissolved iron concentrations in the water column were about 0.2 nM higher than those observed in the other cruises and Martin's data. As dissolved iron concentrations in February 1998 were at low levels similar to those in summer, there is a possibility that iron concentration in winter has some annual or temporal variation. Although deeper surface mixed layer in February 1999 (∼80 m) than in February 1998 (∼50 m) and deposition of aeolian dust from Alaska (Boyd et al. 1998) could have an influence on the surface iron concentrations, these processes cannot explain the increases in dissolved iron at the subsurface layers. Another possible source of iron for deeper water is horizontal transport from the continental margins. Cullen et al. (2009) described that physical mechanisms, such as tidal currents and Ekman transport in the bottom boundary layer, may cause advection of iron-enriched waters from British Columbian continental shelf to offshore regions, while most of the particulate and colloidal iron were found to be removed from the surface waters within the inner- and mid-continental shelf of the northern Gulf of Alaska (Wu et al. 2009). On the other hand, coastal waters in the northern Gulf of Alaska are rich in iron (Lippiatt et al. 2010), and the Alaska stream current could pick up iron-rich coastal water along the Aleutian Island and transport it to Ocean Station P in 4–6 months (Bograd et al. 1999). Lam et al. (2006) provided evidence showing a lateral supply of particulate iron from the continental margin off the Aleutian Islands to Ocean Station P in winter using an ocean general circulation model. However, the influence of coastal water was not clearly supported by salinity and nutrient data in February 1999, and thus, further studies are needed to elucidate the fluctuation of dissolved iron in winter.


Fig. 5. Vertical distributions of the dissolved iron concentrations at Ocean Station P in September 1997 and February 1999. The data reported by Martin et al. (1989) at Station T-7 were also plotted for comparison. Reprinted from Mar. Chem., 74, Nishioka et al., Size-fractionated iron concentrations in the northeast Pacific Ocean: Distribution of soluble and small colloidal iron, 157–179, © 2001, with permission from Elsevier.

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With regard to size-fractionated iron (Fig. 6), labile particulate iron concentration was found to fluctuate at low levels (<0.15 nM) in the water column above 600-m depth. Meanwhile, soluble iron concentrations were low in surface mixed layer, except in February 1999, and were temporally changeable below 100-m depth. Colloidal iron concentrations were also low in the surface, except one depth in September 1997, and showed temporal variation below 100-m depth. From the depth of 200–600 m, colloidal iron represented 13–50% of iron in dissolved fraction. Relatively high concentrations were observed for both soluble and colloidal iron at 200–600-m depth in February 1999, when compared with June and September.


Fig. 6. Temporal change in iron distributions of each size fraction at Ocean Station P in September 1997, June 1998, September 1998, and February 1999. (a) Soluble Fe, (b) colloidal Fe, and (c) labile particulate Fe. Reprinted from Mar. Chem., 74, Nishioka et al., Size-fractionated iron concentrations in the northeast Pacific Ocean: Distribution of soluble and small colloidal iron, 157–179, © 2001, with permission from Elsevier.

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Active decomposition of sinking organic matter was observed at 100–200 m, and good correlation between regenerated nutrient and colloidal iron concentrations was obtained (Fig. 7). These results suggest that the decomposition process of sinking organic matter could be one of the mechanisms for the formation of colloidal iron at the subsurface layer. Relatively low concentrations of iron-complexing organic ligands in the surface water may allow excess dissolved iron to be scavenged efficiently from the surface to the subsurface layer together with the sinking organic particles, and consequently form low dissolved iron concentrations in the surface waters during most of the seasons. On the other hand, sinking flux of organic matter at 200-m depth reaches the maximum in May at Ocean Station P (Wong et al. 1999), and the lowest concentrations of colloidal iron in the deep water have been observed in June. Therefore, it seems that the abundance of colloidal iron in the subsurface layer is also under the control of scavenging process at 200–1000-m depth and would be transported to the deeper layers.


Fig. 7. Comparison of colloidal (200 kDa-0.2 μm) iron with phosphate concentrations at Ocean Station P in September 1997 and February 1999. Reprinted from Mar. Chem., 74, Nishioka et al., Size-fractionated iron concentrations in the northeast Pacific Ocean: Distribution of soluble and small colloidal iron, 157–179, © 2001, with permission from Elsevier.

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In the western subarctic North Pacific, time-series observation of iron distribution was conducted at Station A7 in the Oyashio region and the oceanic subarctic Station B9 (Fig. 8, Nishioka et al. 2007). The vertical profiles clearly showed temporal variability in dissolved and total iron concentrations in the water column at both the stations (Fig. 9). Higher temporal variability of dissolved and total iron concentrations was observed in the entire water column at A7 than at B9. Station A7 is located upstream of the Oyashio and close to the Sea of Okhotsk, while Station B9 is downstream of the flow. Comparison of time-series data between A7 and B9 indicated that some fractions of colloidal and particulate iron were lost from the water column during water transportation. From these spatial and temporal iron distributions, it can be inferred that there is a large source of iron, mainly in the particulate phase, upstream of the Oyashio region and that the iron is distributed to the subarctic water in the western North Pacific.


Fig. 8. Sampling stations for observational studies on iron distributions and for both onboard and in situ iron-enrichment experiments. The arrows indicate a schematic image of the surface water currents.

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Fig. 9. (a) Vertical profiles of dissolved iron (D-Fe), total iron (T-Fe), salinity, and nitrate + nitrite concentrations from January to May at Station A7 (41°30'N, 145°30'E; upstream of Oyashio region) and (b) from March to May at Station B9 (44°N, 155°E; downstream of Oyashio region) in 2003. Nishioka et al., Iron supply to the western subarctic Pacific: importance of iron export from the Sea of Okhotsk, J. Geophys. Res., 112, C10012, © 2007 American Geophysical Union. Reproduced by permission of American Geophysical Union.

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In the Oyashio region, clear seasonal variability of dissolved iron concentrations in the surface mixed layer was observed (Nishioka et al. 2007, 2011). The dissolved iron concentration reached a maximum in January and remained high (average of 0.6 nM) in winter, when deep vertical mixing down to ∼200 m brought up iron-rich subsurface water to the surface (Fig. 10). With the development of the spring blooms under shallower surface mixed layer, the dissolved iron levels decreased to <0.2 nM. Higher concentrations of surface dissolved iron in March (∼0.3 nM), when compared with May (<0.1 nM), were also observed in the oceanic region (B9; Fig. 9). The observed seasonal change in the dissolved iron was similar to that of the nitrate concentration in the surface mixed layer, suggesting biological control of the dissolved iron levels in the surface water during spring.


Fig. 10. Seasonal variations in (a) sea-surface dissolved iron concentrations (average in surface-mixed layer); (b) nitrate + nitrite concentrations (average in surface-mixed layer); (c) surface-mixed layer depths (MLD); and (d) chlorophyll a concentrations (average in surface-mixed layer) from January to the end of May 2003 along the "A-line". Nishioka et al., Iron supply to the western subarctic Pacific: importance of iron export from the Sea of Okhotsk, J. Geophys. Res., 112, C10012, © 2007 American Geophysical Union. Reproduced by permission of American Geophysical Union.

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2-4. Possible sources of iron in the subarctic North Pacific

While regeneration and recycling of biological iron may support the bulk of primary production in the remote oceanic regions with low external iron supply rates (Hutchins et al. 1993, 1995; Hutchins and Bruland 1995; Price and Morel 1998; Sato et al. 2007), it is important to identify sources of new iron for better understanding of the regulation mechanisms of phytoplankton production in the subarctic North Pacific.

In most of the open-ocean regions, the major source of new iron is deposition of wind-borne continental dust, carried for long distances from terrestrial arid regions. The subarctic North Pacific receives most of atmospheric iron input from Asia (Duce and Tindale 1991), while aeolian dust from Alaska is another important source in the eastern subarctic region (Boyd et al. 1998). There is a longitudinal dust gradient across the North Pacific, and the flux of dust containing iron over the western region is an order of magnitude higher than that in the eastern region (Duce and Tindale 1991). However, supply rates largely depend on the prevailing wind patterns and seasonal storms, and delivery is unpredictable and episodic. The main eastward flow of Asian mineral dust passes over the western North Pacific between 30° and 40°N, south of the Subarctic Front, with the mean annual total (dry + wet) deposition flux of 5 g m–2yr–1 (7400 nmol Fe m–2d–1, using a dust iron content of 3%) at 30°–60°N (Uematsu et al. 2003).

Young et al. (1991) have reported on the stimulation of primary production by the dissolution of the mineral aerosol iron associated with the Asian dust in the North Pacific Central Gyre. However, atmospheric deposition is also an important source of nitrogen for primary production in oligotrophic North Pacific Gyre (Duce 1986; Duce et al. 2008). During the above- mentioned dust deposition events, nitrate seems to be deposited along with iron (DiTullio and Laws 1991), and hence, the observed increase in primary production after the dust events reported by Young et al. (1991) could be driven by the simultaneous input of iron and nitrate.

Wide range of total aerosol iron flux (150–4300 nmol Fe m–2d–1, using a deposition velocity of 1 cm s–1) has been reported at the 40°–50°N in the North Pacific in May–June 2002, including some dusty conditions (Buck et al. 2006). On the same cruise, based on the measurements of the dissolved aluminum concentration in the surface water, Measures et al. (2005) concluded that only a small amount of mineral dust enters the surface waters of the western subarctic North Pacific. Bishop et al. (2002) observed increases in carbon biomass in the surface mixed layer after the passage of a cloud of Gobi desert dust in the eastern subarctic North Pacific, but definite evidence for dust-induced increases in primary production and algal biomass has not been observed in the western subarctic North Pacific, probably due to the coexistence of another large iron source (Nishioka et al. 2007).

Iron in the dust particles must first dissolve in the surface water before becoming available to phytoplankton. Solubility of iron from dust had been reported to range from 0.001% to 90%, and this variability seems to be caused by the differences in both aerosol properties and leaching schemes (Boyd et al. 2009). Buck et al. (2006) reported that the solubility of aerosol iron collected over the western subarctic North Pacific (40°–50°N) ranges from 0.3% to 26% in seawater, and their data indicate flux of soluble iron as 6–58 nmol Fe m–2d–1 (using a deposition velocity of 1 cm s–1). On the other hand, the iron solubility of the Asian mineral dust entering the western subarctic North Pacific has been estimated to be ∼0.4% (Ooki et al. 2009). Thus, deposition events of mineral dust during spring (∼0.4 g m–2week–1; Uematsu et al. 2003) will increase the dissolved iron in surface water only by about 0.02 nM, by assuming 3% iron content and 50-m deep surface mixed layer. In fact, Boyd et al. (2009) pointed out that dust-mediated phytoplankton blooms are probably rare in the modern ocean due to the slow dissolution of iron in the surface seawater. However, a recent study by Iwamoto et al. (2011) reported that sea fog can increase the air-to-sea transfer of Asian dust and supply a dissolved iron of 360–5900 nmol Fe m–2 event–1 in the northwestern (NW) North Pacific near Japan. If such large atmospheric iron deposition occurs every spring months, some increase in the surface dissolved iron concentrations should be detected; however, multiyear (2003–2008) time-series observation of the dissolved iron in the Oyashio region showed that surface dissolved iron is mainly controlled by upward transport of iron-rich intermediate water by tidal and winter mixing processes as well as biological uptake during the spring bloom period. They concluded that dust inputs are not the major driving phenomenon of the annual iron cycle in the Oyashio region, but instead, sporadic dust iron input may have a role for sporadically occurring phytoplankton blooms, although such impact would be spatiotemporally limited. At present, we could not obtain an appropriate estimate for iron released from the deposited dust at the initial contact with seawater as well as the following dissolution processes during its mixed layer residence time, and therefore, the impact of dust iron on phytoplankton growth in the subarctic Pacific is still not clear.

Although iron from mineral aerosols dominates most of the ocean deposition, iron produced during combustion, both from industrial and biomass burning sources, may be important because iron solubility of the combustion source (4.0%) is ten times higher than that of the mineral dust (Luo et al. 2008). By using the Dust Entrainment and Deposition model, Luo et al. (2008) estimated that deposition of combustion iron at the western subarctic ranges 0.05–0.2 ng Fe m–2s–1 (80–300 nmol Fe m–2d–1), which corresponds to 5–10% of the total iron deposition, and that deposition of soluble iron from combustion represents >20% of the soluble iron deposition over much of the open ocean. This implies that shifts in the quantity of soluble iron associated with industrial activity in Asia may play a role in the changes in the iron availability for phytoplankton in the North Pacific.

Another interesting atmospheric iron source is volcanic ash, which may also have an important role in the injection of biologically available iron into the surface ocean (Duggen et al. 2010). Recently, Hamme et al. (2010) demonstrated that volcanic eruption in the Aleutian Islands had spread volcanic ash over much of the subarctic NE Pacific in August 2008 and initiated one of the largest phytoplankton (diatom) blooms observed in this region. In the summer of 2008, three volcanoes in the Aleutian Islands and two volcanoes in the Kamchatka Peninsula released unusually large amount of volcanic ash that was deposited in the North Pacific. We also observed a blooming of diatoms 4–5 days after a volcanic ash deposition event in the western subarctic North Pacific in August 2008 and found that processing of ash particles in the acidic fog increased both the deposition velocity of ash particles and its iron solubility (Takeda et al. in preparation). These observations clearly provide pieces of evidence that subarctic phytoplankton has a potential to respond to atmospheric iron supply if it occurs at an optimum time of the year. Volcanic eruptions occur occasionally in the Aleutian and Kamchatka regions, and thus, it would be important to watch the activities of these volcanoes and the possible impacts on phytoplankton productivity in the subarctic North Pacific.

In addition to aeolian dust inputs, substantial iron comes from subsurface waters that are enriched with regenerated iron from sinking organic particles by deep vertical mixing during winter. The high iron concentrations observed at the subsurface layer in the western subarctic region suggest that the subsurface iron is an important iron source for phytoplankton in early spring, although this source is not available during stratified summer. Contribution of winter mixing in the annual upward iron flux has been estimated to be 59% in the Oyashio region and 32% in the western subarctic Pacific, and the total upward fluxes in the western region (11–13 μmol Fe m–2yr–1 or 31–35 nmol Fe m–2d–1 ) is about 4 times higher than that in the eastern subarctic Pacific (Nishioka et al. 2007). As soluble iron flux by dust deposition in the western North Pacific is around 30 nmol Fe m–2d–1 (using 7400 nmol Fe m–2d–1 and 0.4% dissolution), the annual upward iron flux is comparable to the atmospheric iron flux. Iron supplied from below usually depletes before the complete drawdown of macronutrients because of the lower ratio of iron to macronutrients in the subsurface water than the ratios required by phytoplankton. The eastern subarctic Pacific water contains only 4 pM Fe (μM NO3)–1 in the subsurface gradient, while Oyashio and western subarctic Pacific waters have significantly higher ratios (44–52 pM Fe (μM NO3)–1; Nishioka et al. 2007).

As one of the sources of iron in the intermediate water of the western subarctic North Pacific, Nishioka et al. (2007) pointed out the importance of lateral transportation of iron-rich intermediate waters that contain resuspended iron from the continental shelf areas of the Sea of Okhotsk to a wide area of the western subarctic North Pacific (Fig. 11). This source of iron, mainly in the particulate phase, is supplied to the surface layer by diffusion and deep winter mixing in the Oyashio region as well as diapicnal mixing at the Kuril Straits. Lam and Bishop (2008) also suggested lateral advection of labile dissolved and particulate iron from the continental shelf as well as the upper continental slope of the Kuril/Kamchatka margin as an important source for high particulate iron in the western subarctic Pacific.


Fig. 11. Schematic of the iron supply process in the subarctic NW Pacific. Water ventilation processes in this region control the transport of dissolved and particulate iron through the intermediate water layer from the continental shelf of the Sea of Okhotsk to the wide area of the western subarctic Pacific. Nishioka et al., Iron supply to the western subarctic Pacific: importance of iron export from the Sea of Okhotsk, J. Geophys. Res., 112, C10012, © 2007 American Geophysical Union. Reproduced by permission of American Geophysical Union.

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A longitudinal section of dissolved and total iron along 165°E showed that high iron concentrations extend southward at low-salinity intermediate water depth, which had been influenced by North Pacific Intermediate Water (NPIW) formation in the subarctic (Fig. 12). The high iron concentration in the intermediate water was found to be derived from the resuspended sediments of the Sea of Okhotsk and coastal sediments of the Kuril Islands (Nishioka et al. 2007). A high total iron core in the intermediate layer at 35°N was observed to correspond to the eastward-turned Oyashio flow subducted below the warm Kuroshio extension water mass. The time scale of the NPIW transportation from the subarctic region to subtropical area has been estimated to be a few decades (Watanabe et al. 1994). Due to the loss of iron during water transport, the available iron is not sufficient for complete utilization of upwelled nitrate in the HNLC region of the western subarctic North Pacific.


Fig. 12. A longitudinal section of dissolved iron and total iron profile in the North Pacific along 165°E. A density range of 26.6–27.5σθ is located on each figure. Nishioka et al., Iron supply to the western subarctic Pacific: importance of iron export from the Sea of Okhotsk, J. Geophys. Res., 112, C10012, © 2007 American Geophysical Union. Reproduced by permission of American Geophysical Union.

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In the eastern subarctic North Pacific, the surrounding continental margin provides a subsurface supply of iron to the interior of the eastern subarctic Pacific (Johnson et al. 2005; Lam et al. 2006; Cullen et al. 2009). Although there is an observation of the increase in carbon biomass after the passage of Asian dust storm in spring (Bishop et al. 2002) as well as episodic dust supply from the Alaskan source regions (Boyd et al. 1998) and volcanic ash deposition in summer (Hamme et al. 2010), almost constantly low algal biomass and production, observed by time-series studies at Ocean Station P, suggests that such atmospheric iron supply is extremely rare in the eastern subarctic Pacific (Frost 1991; Wong et al. 2002). However, in coastal regions, there is high iron supply from riverine input, resuspension of bottom sediments, and vertical mixing of iron-rich deeper water (Martin et al. 1989; Nishioka et al. 2001; Chase et al. 2005; Wu et al. 2009; Lippiatt et al. 2010). Iron from these coastal and shelf sources can be delivered to the middle of open ocean by mesoscale eddy transport, tidal currents, and Ekman transport from the Canadian coast to the western oceanic region (Johnson et al. 2005; Crawford et al. 2007; Cullen et al. 2009) and southeastward advection from the Aleutian continental shelf via the Alaska Gyre (Lam et al. 2006).

The anticyclonic mesoscale Haida eddies form off the west coast of Canada in winter and then track roughly westward into the open ocean (Crawford and Whitney 1999). These eddies carry large quantities of iron-, macronutrient-, and chlorophyll-rich coastal waters, and some eddies transport them into the HNLC waters of the Alaska Gyre, although most eddies stay relatively near the shelf/slope regions. Iron concentrations in the surface mixed layers of the eddies decrease to low levels, similar to those of outside waters within 4 months, but the subsurface eddy core waters have been found to contain 1.5–2 times more iron than the surrounding waters even 16 months after its formation (Johnson et al. 2005). The steady vertical iron transport from the eddy core waters into the photic zone would occur throughout the lifetime of the eddy, and it has been considered to make a significant contribution to the biologically available iron in this region. A recent study also suggested that eddies in the Alaskan Stream along the Aleutian Islands could have a significant impact on the nutrient and biota exchange between the coastal area south of the Aleutian Islands and the offshore region in the western and central subarctic North Pacific, similar to the eddies in the Gulf of Alaska (Ueno et al. 2009).

The depth of the continental shelf surrounding the Gulf of Alaska is close to the depth of the year-round pycnocline (∼150 m) in the northeast subarctic Pacific (Whitney and Freeland 1999). Thus, lateral transport by tidal current along the constant density surface would carry iron from the continental shelf to the open ocean (Cullen et al. 2009). Lam et al. (2006) found iron from the continental margins in the suspended particles collected from the surface mixed layer at Ocean Station P in winter. Increase in labile particulate iron was also observed there in February 1999, together with high soluble iron (Nishioka et al. 2001) The Alaska Stream current could pick up iron-rich coastal water along the Aleutian Island and a strong recirculation from the head of the Gulf in winter is likely to transport it to Ocean Station P in 4–6 months (Bograd et al. 1999). This subsurface delivery of iron from continental shelves and deep wintertime mixing could be a possible iron source for phytoplankton in winter, but further studies are needed to estimate the supply rate of particulate iron from the sediment, its residence time in the subsurface/surface layers, and the release rate of biologically available dissolved iron from the particles.

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3. Iron requirement for growth of phytoplankton

Extensive differences among phytoplankton in their adaptations to iron availability have been identified through laboratory studies. Brand et al. (1983) and Brand (1991a) observed that growth rates of oceanic phytoplankton are less limited by low concentrations of iron than those of coastal species. However, cellular iron uptake rates are similar among the neritic and oceanic species when rates are normalized to cell surface area, and the ability of the oceanic species to outgrow its neritic congener at low iron concentrations is almost entirely due to its low cellular iron requirement (Sunda et al. 1991; Sunda and Huntsman 1995). Laboratory studies have also shown that sub-optimum iron levels probably restrict the development of large phytoplankton cells more than small cells (Hudson and Morel 1990). These differences indicate that iron is an important environmental factor. Therefore, changes in the rate of iron input and recycling in nutrient-replete regions of the subarctic North Pacific may result in species and cell-size shifts in phytoplankton assemblage. However, data on the iron demand for the growth of the North Pacific phytoplankton species are limited (Marchetti et al. 2006a). In addition to the importance of studying representative species of the North Pacific phytoplankton community, it is important to use local isolates because of the genetic differentiation between the populations (Brand 1991b).

Phytoplankton growth rate, μ, is typically described according to a Monod saturation function of the form μ = μmax[S/(Ks+S)], where S is the limiting nutrient concentration, μmax is the maximum nutrient-saturated growth rate, and Ks is the half-saturation constant for growth with respect to S. The value Ks is of particular importance in numerical ecosystem models because it is used to determine the nutrient that is limiting at any given time, by comparison of the ambient nutrient concentration with Ks (e.g., Takeda et al. 2006). In the case of iron, only a few estimates of Ks have been reported based on field measurements (Noiri et al. 2005; Kudo et al. 2006).

In this section, the differences among species in their growth rates are described as a function of dissolved iron concentration using laboratory cultures of phytoplankton species isolated from the western North Pacific, and the results of shipboard experiments conducted to identify the concentrations in which iron limits the growth of natural phytoplankton assemblage in the eastern and western subarctic North Pacific are discussed to understand the dominance of some phytoplankton species/groups under certain iron conditions.

3-1. Laboratory and onboard experimental methods for estimation of phytoplankton iron requirement

Growth rates were measured as a function of dissolved iron concentrations for two clones of haptophytes and three clones of centric diatoms isolated from the western North Pacific (Table 2). Although sterile techniques were used and bacteria were never apparent in the cultures, most were probably not axenic. The experimental culture media were prepared from filtered (0.03-μm pore-size) western North Pacific (39°N; 155°E) surface water enriched with nutrients (30 μM NaNO3, 1 μM NaH2PO4, and 25 μM Na2SiO3), trace metals (10 nM ZnCl2, 10 nM MnCl2, 1 nM (NH4)6Mo7O24 , and 1 nM CoCl2), and vitamins (0.04 nM Vitamin B12, 0.2 nM Biotin, and 30 nM Thiamin HCl). The major nutrient stock solutions and the filtered surface water were treated with Chelex-100 resin to remove trace-metal contaminants (Morel et al. 1979; Price et al. 1988/1989). Various levels of FeEDTA were added: none, 0.1, 1, and 10 nM. The final concentration of Na2EDTA was adjusted to 10 nM in all treatments to avoid the problem of overchelation at low iron concentrations, which could prevent the phytoplankton from taking up most of the iron, and to minimize the precipitation of iron to a form unavailable to phytoplankton (Brand 1991a). Although iron uptake by phytoplankton is typically correlated with free hydrated Fe3+ in seawater containing high concentrations (∼100 μM) of synthetic chelators, it is the labile hydrolysis species comprising Fe(III)' and Fe(II)' that actually control the uptake rates (Wells et al. 1995). The total iron was computed from the sum of the concentration of added FeEDTA and the background iron concentration. Background iron in the medium was measured as 0.04 nM by chelating resin concentration and chemiluminescence detection methods (Obata et al. 1993). The media were sterilized by filtration using acid-washed 0.2-μm Teflon filter. The cells were grown in triplicate acid-washed microwave-sterilized 25 × 90-mm polycarbonate tubes containing 20 ml of medium at 20 ± 0.5°C. A light intensity of 170 μmol m–2s–1 was provided by white fluorescent bulbs with a 14L:10D photoperiod. This light intensity surpasses the light saturation level of natural phytoplankton assemblages in the western subarctic North Pacific in summer (100 μmol m–2s–1, Noiri et al. 2005). Experimental cultures were grown for 4 times in continuous batch culture as sequential replicates and monitored by measuring in vivo fluorescence with a Turner 10-AU-005 fluorometer (Brand et al. 1983). The specific growth rate of each sequential replicates was determined by the least-squares method of linear regression on the logarithmically transformed data. The maintenance of continuous batch cultures allowed for the determination of acclimated growth rates, although reproduction was not completely constant for diatoms because of the changes in cell size (Brand et al. 1981).


Table 2. Phytoplankton strains isolated from the NW Pacific for laboratory culture experiments.

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An iron-enrichment bottle incubation experiment was performed to determine the relationship between iron concentration and phytoplankton growth in HNLC water of the eastern subarctic North Pacific. Surface water sample with its resident phytoplankton population was collected from 20-m depth at Ocean Station P (50°N; 145°W) using trace metal clean Teflon bellows pump system during the cruise 9829 of CCGS J. P. Tully in September 1998. The water samples dispensed from a 20-L polycarbonate carboy into 250-ml polycarbonate bottles were spiked with 0, 0.025, 0.05, 0.1, 0.25, 0.5, 0.75, 1.0, 2.5, or 5 nM iron (as FeCl3) and then incubated in on-deck incubators at surface-water temperature under 30% light level. The water samples were not prescreened with netting to exclude grazing organisms from the incubation bottles. Changes in the concentrations of chlorophyll-a (size fractionation at 5 μm) and nutrients were monitored on Days 1, 2, 3, 4, 5, and 7 to calculate the phytoplankton growth rate. Twelve bottles were used for each treatment, two bottles were withdrawn from the incubator at a time, and the incubation bottles were not repetitively sampled to avoid contamination problem. Initial samples were drawn directly from the 20-L polycarbonate carboy. Trace-metal clean techniques for bottle cleaning and sample treatment were applied according to the protocol described by Takeda and Obata (1995). As ambient seawater could have subnanomolar iron as a background, it would be efficient if we could reduce the biologically available iron level in the incubation bottle to near zero by adding excess strong iron-complexing ligands. A fungal siderophore, deferoxamine mesylate (DFB), was used for this purpose, and its effect was examined by adding 100 nM DFB with 1 nM iron simultaneously, and another without iron.

During the second in situ iron-enrichment experiment in the western subarctic Pacific, the Subarctic Pacific Iron Enrichment for Ecosystem Dynamics Study II (SEEDS-II) (Tsuda et al. 2007), an onboard bottle incubation experiment was conducted 17 days after the first iron infusion to examine the effect of iron addition on net growth rate of the iron-induced bloom at the decline phase. Surface water was collected from 10-m depth at the center of the iron-fertilized patch water on Day 17 by using acid-cleaned Teflon-coated 10-L X-Niskin bottles suspended on a Kevlar hydro-wire during the KM0415 cruise of R/V Kilo Moana. The seawater was homogenized in an acid-cleaned 25-L polycarbonate carboy and then dispensed into 250-ml polycarbonate bottles. The water samples were spiked with 0, 0.5, 1.0, 2.0, or 5 nM iron as FeCl3 and incubated in on-deck incubators at surface water temperature under 30% light level. The water samples were prescreened with 202-μm Teflon mesh netting to exclude grazing organisms from incubation bottles. Triplicate bottles were withdrawn from the incubator after 4 and 6 days and submitted to the measurements of chlorophyll-a (size fractionation at 5 μm) and nutrients.

3-2. Growth responses of laboratory culture strains isolated from the North Pacific

The phytoplankton species investigated in this study were found to differ in their growth response to iron concentrations (Fig. 13). The maximum growth rates of two haptophyte clones were observed at iron concentrations over 1 nM, and these clones showed moderate growth, which corresponded to half of the maximum rate, even at the lowest iron concentration. Within the iron concentrations examined, the growth rates of Emiliania huxleyi were always higher than those of Gephyrocapsa oceanica. Large diatoms, Actinocyclus sp. and Thalassiosira sp., could not grow well at 0.04 and 0.14 nM iron, but their growth rates increased to 0.75–1.1 d–1 at >1 nM iron. The growth rate of Chaetoceros sp. increased consistently with the increase in the iron concentration from 0.4 to 1 d–1, indicating the subsistence optimum iron concentration above 10 nM.


Fig. 13. Relationships between total dissolved iron concentration and specific growth rate (a, b) and final cell yield measured at the stationary phase (c, d) for Emiliania huxleyi (○), Gephyrocapsa oceanica (■), Thalassiosira sp. (▲), Actinocyclus sp. (□), and Chaetoceros sp. (●) isolated from NW Pacific.

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The plots of specific growth rates vs. iron concentration are well described by the Monod saturation function (not shown). By assuming that the iron concentration in each experimental treatment during the exponential growth phase was close to the initial value, linear regressions of S/μ against S (Woolf plot) using a least-squares method yielded the Ks estimates for iron. The Ks values for Emiliania huxleyi and Gephyrocapsa oceanica were 0.03 nM, much lower than those of Chaetoceros sp. (0.20 nM), Thalassiosira sp. (1.6 nM), and Actinocyclus sp. (1.1 nM).

The final cell yields of each phytoplankton species counted at the stationary phase are shown in Fig. 13. Both the final cell yield and the growth rate showed similar increasing pattern with the increase in iron concentration, but the difference between minimum and maximum cell yields was only 3, 6, and 17 times for E. huxleyi, G. oceanica, and Chaetoceros sp., respectively. By assuming that almost all the iron in the medium was used up by the iron-starved cells at the stationary phase under the iron-limited growth conditions (Brand 1991a) and that adsorption of added iron to the wall of acid-cleaned polycarbonate tubes was negligible (Takeda and Obata 1995), cellular Fe:C ratios were estimated for each clones by using particulate organic carbon (POC) values measured at the stationary phase (Table 3). Two haptophyte clones had low Fe:C ratios of <1 μmol Fe:mol C at the lowest iron condition, and similar ratios were observed at 0.14 nM iron. On the other hand, Fe:C ratio of Chaetoceros sp. at 0.04 nM iron was 2.5 times lower than that at 0.14 nM iron.


Table 3. Cellular iron to carbon ratio (Fe/C) and iron use efficiency of the NE Pacific phytoplankton under iron stress. Estimation was based on the measured POC value at the stationary phase of the growth and by assuming that all the iron was taken up under the iron-limited condition.

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The observed differences among the species in their growth responses to the iron concentration suggest that iron can be one of the important environmental factors controlling phytoplankton growth rate and species composition of phytoplankton in the North Pacific, where dissolved iron concentrations in the surface water range at subnanomolar level. Small cells, such as Chaetoceros sp., E. huxleyi, and G. oceanica, showed low Ks values and low Fe:C ratios, indicating high affinity for iron uptake as well as adaptation to low ambient iron concentrations by reducing minimum iron cellular quotas. Marchetti et al. (2006a) reported Fe:C ratios for four oceanic Pseudo-nitzschia spp. in a range of 2.8–3.7 μmol Fe mol C–1 under low iron conditions. These values are similar to those of Chaetoceros sp. estimated in this study. The ratios of E. huxleyi and G. oceanica were one-order lower than the Fe:C ratio (3.1 μmol Fe mol C–1) obtained for E. huxleyi at relatively high growth rate (0.88 d–1, Sunda and Huntsman 1995). Previous culture studies also showed that oceanic coccolithophores have particularly low iron requirements (Brand et al. 1983; Brand 1991a). In addition, E. huxleyi has been reported to take up iron at a faster rate per unit area of cell surface than diatoms and dinoflagellates (Sunda and Huntsman 1995). The estimated Fe:C ratios for E. huxleyi and G. oceanica were found to be lower than the ratio reported for oceanic diatom Thalassiosira oceanica (2.5∼5.4 μmol Fe mol C–1, Sunda et al. 1991; Sunda and Huntsman 1995; Maldonado and Price 1996; Marchetti et al. 2006a), which is known to have greatly reduced its iron requirements for growth by decreasing the cellular concentrations of cytochrome β6/f and PSI (Strzepek and Harrison 2004), by synthesizing flavodoxin instead of ferredoxin (Strzepek and Harrison 2004), and by using copper-containing plastocyanin instead of cytochrome c6 for electron transport (Peers and Price 2006). Their moderate growth rates (0.44–0.56 d–1) even at 0.04 nM iron suggest their high ability to survive under the iron-limited open-ocean environment, although chemical speciation of iron in natural seawater containing natural-iron-complexing organic ligands could be different from the culture media in which labile inorganic iron species were controlled by the synthetic chelator, EDTA.

Large diatoms, such as Actinocyclus sp. and Thalassiosira sp., had high Ks of >1 nM and high Fe:C ratios similar to the coastal species (Sunda et al. 1991; Sunda and Huntsman 1995; Marchetti et al. 2006a). The ability of the cells to grow at low iron concentrations can be determined by the iron-use efficiency (growth rate per mole of cellular iron, mol C [mol Fe]–1 d–1). These values of E. huxleyi and G. oceanica were found to be 3∼10 times higher than the value of Chaetoceros sp. (Table 3). The iron-use efficiency of Thalassiosira sp. was observed to be 5 times higher than that of Actinocyclus sp., which had a value similar to that of Chaetoceros sp. and the oceanic Pseudo-nitzschia spp. (2.1∼3.6 × 105 mol C [mol Fe]–1 d–1, Marchetti et al. 2006a). The Oyashio waters in the western subarctic North Pacific are known to have a dissolved iron concentration of more than 1 nM in winter (Nishioka et al. 2007), and this iron-rich water could support moderate growth of large diatoms, such as Actinocyclus sp. and Thalassiosira sp., during the spring bloom. The obtained experimental results are consistent with the fact that growth of large diatoms, mainly Thalassiosira spp., was enhanced upon addition of 1 nM iron in the onboard incubation experiment carried out in the western North Pacific (see Subsection 4.2). On the other hand, the lower Fe:C ratio for oceanic Pseudo-nitzschia spp. (Marchetti et al. 2006a), when compared with the large centric diatoms, may explain the observed increase in cell density of pennate diatoms in the incubation experiments conducted at Ocean Station P (see Subsection 4.4), where the surface dissolved iron concentrations are usually <0.1 nM. Such bottle incubation experiments using natural plankton assemblage indicate that the addition of iron caused shifts in the community structure from nanoplankton-dominated population to one dominated by micro-size diatoms. On the other hand, results from the laboratory culture showed that the growth rates of small haptophytes at 1 nM iron were similarly high to those of the micro-size diatoms (Fig. 13). This suggests that the growth of nano-size phytoplankton in the iron-enriched bottles was inhibited by other rate- limiting factors. In particular, microzooplankton grazing appears to be the dominant process in maintaining the nano-size phytoplankton population at a low level (Welschmeyer et al. 1991).

The estimated Ks values of western North Pacific phytoplankton were compared with those reported for oceanic strains isolated from the Atlantic and Pacific Oceans (Table 4). However, it is difficult to directly compare each data due to the differences in the iron concentration used for the Ks estimation. In many of the other laboratory experiments, high concentrations of EDTA (100–500 μM) were added together with high concentrations of total iron (1–10000 nM) to regulate the free-ion concentration of iron at low levels (Sunda and Huntsman 1995; Timmermans et al. 2001), and thus, the concentration of dissolved inorganic iron [Fe'] was used instead of total dissolved iron for the estimation of Ks, which resulted in values that were one or two orders of magnitude lower than those obtained in this study. In the buffered culture media containing high concentration of EDTA, the concept of bioavailability should be tied to kinetics because iron-EDTA complex dissociates over a finite timescale and limitation could occur if the dissociation rate is too slow to allow the phytoplankton to grow. Despite these disparity, the Ks obtained in this study are similar to the value averaged for 12 diatom species (0.35 ± 0.44 nM), which includes not only the subarctic and subtropical oceanic isolates but also the coastal and Antarctic strains (Sarthou et al. 2005). However, some organically complexed iron may be a source of available iron for phytoplankton growth (Hutchins et al. 1999b; Maldonado and Price 1999), and phytoplankton may release iron-complexing ligands into the culture media (Boyé and van den Berg 2000). In addition, >99% of the dissolved iron is organically complexed in natural seawater (Gledhill and van den Berg 1994; Rue and Bruland 1995). Therefore, more studies are needed to know which iron species is/are available for phytoplankton in seawater and culture media before the application of the laboratory experimental results to natural ecosystems.


Table 4. Comparison of the half saturation constant for iron (Ks) for the NW Pacific phytoplankton with the reported value for other oceanic strains.

*The concentration of dissolved inorganic iron (Fe') was used instead of the total dissolved iron for the estimation of Ks. (a) Timmermans et al. (2001); (b) Sunda and Huntsman (1995); (c) Kudo and Harrison (1997).

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3-3. Growth responses of natural assemblage to the changing iron concentrations in the subarctic Pacific

To examine the effect of iron addition on the growth of natural phytoplankton assemblage and to estimate a threshold of iron limitation, iron-enrichment bottle incubation experiments were performed at Ocean Station P in September 1998. The surface water used for the incubation experiments was characterized by low phytoplankton biomass (0.31 μg L–1 chlorophyll-a) and low concentration of dissolved iron (0.14 nM). High concentrations of nitrate (6.3 μM), phosphate (0.85 μM), and silicic acid (11 μM) were observed at the end of summer, indicating possible iron-limitation stress of the ambient phytoplankton assemblage. Phytoplankton of size <5 μm contributed to nearly half of the total chlorophyll-a biomass.

Figure 14 shows the time variations in chlorophyll-a and the nutrients in the incubation bottles for the various added concentrations of iron. Chlorophyll accumulation and nutrient consumptions in the incubation bottles generally increased during the course of the experiments, in amounts related to the concentrations of iron added to each treatment. At the end of the 7-day incubation period, chlorophyll-a concentrations of 0.7–5 μm size-fraction showed similar values when iron was added at 0.5–5.0 nM. On the other hand, there was clear enhancement in the final chlorophyll-a concentration of >5 μm size-fraction between 1.0 and 2.5 nM iron additions. Decreases in silicic acid concentration were observed during the later phase of the experiments in the bottles enriched with iron of 0.25 nM or more, which suggests stimulation of diatoms growth under these iron concentrations with some lag periods.


Fig. 14. Changes in the concentrations of size-fractionated chlorophyll a (a, 0.7–5 μm; b, >5 μm), silicic acid (c), phosphate (d), and nitrate (e) in control and experimental bottles enriched with various amounts of iron. The experiments were conducted at Ocean Station P (50°N; 145°W) in September 1998.

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These data indicate that phytoplankton community growth rates increased with added iron both in small (0.7–5 μm) and large (>5 μm) cells. The net specific growth rates (d–1) of small- and large-size fractions were estimated for each experimental treatment based on the change in chlorophyll-a concentration. A chlorophyll-based estimate of growth may sometimes result in overestimation of the rate due to the increase in cellular chlorophyll-a content with the addition of iron (Takeda and Kamatani 1989). In these experiments, there was a significant correlation between chlorophyll-a and nitrate concentrations in the incubation bottles (r2 = 0.94), and hence, phytoplankton biomass, as measured by chlorophyll-a, was used as a proxy in this study.

While the results showed that chlorophyll-a increases with the increasing dissolved iron concentration in both the size classes, a difference in terms of the response of their growth rate to the addition of iron was observed. Figure 15 shows the relationship between the specific growth rate and the initial dissolved iron concentration (i.e., ambient + added iron) for small (0.7–5 μm) and large (>5 μm) phytoplankton cells. It should be noted that direct comparison between Figs. 13, 15 (and 16b) is not appropriate because of the difference in the major chemical forms of the dissolved iron. The small cells had higher growth rates at iron concentrations <1 nM, but the large cells grew faster than the small cells under high iron conditions. This is not surprising because large phytoplankton are particularly sensitive to iron limitation owing to the relatively low surface area to volume ratio, which decreases the diffusion rate of iron from the bulk solution and the availability of membrane area to anchor necessary transporters.


Fig. 15. Specific growth rates of phytoplankton assemblage, calculated using concentrations of chlorophyll a in two size classes (0.7–5 μm [] and >5 μm []), plotted against initial dissolved iron concentrations. The data were from the bottle incubation experiments at Ocean Station P (50°N; 145°W) in September 1998. The open triangles (0.7–5 μm [] and >5 μm []) represent data from the incubation bottles treated with 100 nM desferrioxamine B.

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Specific growth rates of phytoplankton in the experiments enriched with strong iron-complexing ligands, such as DFB, are also plotted in Fig. 15. Existence of 100 nM DFB reduced the growth rates down to <0.1 d–1 in both the size classes. This finding is consistent with the other experimental results demonstrating that the addition of excess concentration of DFB essentially eliminated iron uptake in natural phytoplankton populations (Wells 1999; Hutchins et al. 1999a). However, the growth rate of large cells recovered to a level, as observed in the control bottles, when 1 nM iron was added together with strong iron-complexing organic ligand, such as DFB. DFB has strong specificity and high conditional stability constant for Fe(III), and it shows very rapid complexation kinetics (Hudson et al. 1992). Thus, the observed slight, but continuous, growth of large-size phytoplankton might indicate their ability to utilize organically complexed iron in seawater, as growth of larger cells using stored intracellular iron could occur only at the initial few days. Maldonado et al. (2006) showed pieces of evidence that iron-limited oceanic diatom, Thalassiosira oceanica, can take up iron from the DFB-iron complex by using copper-dependent high-affinity iron transport system.

The iron-response curves were derived from a nonlinear curve fitting of the data to the Monod equation by assuming that the initial dissolved iron concentration did not change significantly during the incubation. The half-saturation constant, Ks, varied from 0.34 nM for the small cells to 0.53 nM for the large cells, and the maximum growth rate of the large cells (0.73 d–1) was 13% higher than that of the small cells (0.65 d–1). The observed results confirm the importance of cell size in determining iron requirement and the growth response to changing iron concentrations (Morel et al. 1991). The ambient dissolved iron concentration (0.14 nM) was at least two times lower than the estimated Ks values, indicating that iron may be the major limiting factor controlling the community growth rate of in situ phytoplankton.

The Ks values estimated from the experiment at Ocean Station P are within the same order of magnitude as other Ks estimates reported for the phytoplankton community in the subarctic North Pacific (Table 5). Although the Ks values (<0.1 nM) reported by Kudo et al. (2006) are in ranges similar to those of Antarctic phytoplankton communities (Blain et al. 2002; Coale et al. 2003), these low values seem to be partly derived from significant growth of nano- and microphytoplankton (0.3 d–1) in the control bottles, and such changes in chlorophyll and nutrients in the control bottles would be problematic for the estimation of Ks values. Furthermore, their experiments were conducted at 90% of surface irradiance, which is 2–4 times stronger than that at the depth of sample-water collection. Higher irradiance may stimulate algal photosynthetic activity and could also have an effect on iron availability by enhancing photochemical redox cycles. Although relatively low Ks values had been estimated for phytoplankton community during the Subarctic Ecosystem Response to Iron Enrichment Study (SERIES) experiment (Kudo et al. 2006), these results suggest that not only the large species, such as diatoms, but also small algal species are substantially limited by iron availability both in western and eastern subarctic North Pacific. Considering the differences in the biological and geochemical characteristics between the Western Subarctic Gyre and the Alaska Gyre (Harrison et al. 1999, 2004; Suzuki et al. 2002), it is surprising to observe similar Ks values in these two gyres, as algal species composition as well as other environmental factors, such as light and availability of other nutrient elements, are likely to influence the iron requirement of the phytoplankton community. For better understanding of the ecological role of iron limitation in the subarctic North Pacific, further studies are needed on the relative sensitivity of individual taxa from natural waters to iron supply at sufficient low levels that are often observed in the natural environments.


Table 5. Comparison of the half saturation constant for iron (Ks) for subarctic phytoplankton assemblage in two size classes.

*The total (unfiltered) acid-labile iron concentration in the control bottles was measured on the last day of incubation. (a) Kudo et al. (2006); (b) Noiri et al. (2005). Iron was added as a ferric chloride solution in these experiments as well as in this study.

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In addition to the differences in algal species composition and environmental factors, physiological history of phytoplankton could also have important influences on their iron requirement, as luxury uptake (Sunda and Huntsman 1995) and storage of iron (Marchetti et al. 2009) have been reported for certain phytoplankton species. During the second in situ iron-enrichment experiment in the western subarctic Pacific, SEEDS-II, the growth response of phytoplankton at the decline phase of the iron-induced bloom was examined by a series of iron-addition bottle incubation experiments. In SEEDS-II, the surface chlorophyll-a concentration in the iron-enriched patch water increased from Day 4, peaked on Days 11–13, then gradually decreased after Day 13, and finally returned to the initial level on Day 20 (Tsuda et al. 2007; see Section 5). The dissolved iron concentration in the iron-enriched patch increased from the initial 0.02 nM to 1.38 and 0.66 nM after the first and second iron infusions, respectively (Nishioka et al. 2009).

The bottle incubation experiments were conducted on Day 17, when dissolved iron concentration in the iron-enriched patch decreased to 0.07 nM. The ambient phytoplankton community was a mixture of several algal groups, and the contribution of diatoms was 16% of the total chlorophyll-a biomass on Day 17 (Suzuki et al. 2009). The addition of iron largely increased chlorophyll-a concentration, mainly in large-size fraction (>10 μm) within 6 days of the incubation period (Fig. 16a). Large decreases in nitrate and silicic acid concentrations with increasing iron concentration suggest that diatoms became dominant in the iron-enriched bottles.


Fig. 16. (a) Changes in concentrations of size-fractionated chlorophyll a in control and experimental bottles enriched with various amounts of iron, measured at the beginning of the experiment (initial) and after 4 and 6 days of incubation. (b) Relationships between initial dissolved iron concentration and specific growth rates of phytoplankton assemblage, calculated using size-fractionated chlorophyll a data. The onboard incubation was conducted during the SEEDS-II experiment using iron-fertilized surface water on Day 17.

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The net specific growth rates of the small (0.7–10 μm) and large (>10 μm) phytoplankton were estimated based on the changes in chlorophyll-a concentration between 4 and 6 days of the incubation period. As macronutrient concentrations were not depleted at the end of the incubation, chlorophyll-a was used as a proxy for phytoplankton biomass, and the exponential growth was assumed. The relationship between growth rate and iron concentration suggests that more than 2 nM dissolved iron was needed to saturate the iron requirement for growth in both small- and large-size phytoplankton groups (Fig. 16b). The moderate growth rate of large phytoplankton observed in the control bottles, which contained only 0.07 nM dissolved iron as a background, could be explained by their ability of taking up and maintaining excess iron after the iron infusions. However, high iron requirement of phytoplankton community seems to reflect low biological availability of added iron in the incubation bottles.

During the decline phase of the iron-induced phytoplankton bloom in SEEDS-II, large increase in iron-complexing organic ligands was observed with the concentration as high as 2.12 nM on Day 17 (Kondo et al. 2008). It is possible that iron-complexing organic ligands are released to surface water by active zooplankton grazing on phytoplankton (Sato et al. 2007), siderophore production by iron-limited bacteria, and phytoplankton cell lysis by virus infection during the bloom decline phase. Therefore, the iron added to the incubation bottle could have been complexed with the released organic ligands, and consequently, its biological availability could have been greatly reduced. These findings clearly show the importance of chemical speciation of iron in analyzing the growth response of phytoplankton to iron supply, as well as the strong needs for elucidating the form of iron that is available for phytoplankton growth.

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4. Effects of iron enrichment on the phytoplankton growth and nutrient utilization

After the pioneering works by John Martin, several bottle enrichment experiments have been performed in the HNLC waters in the subarctic North Pacific. Experiments in the subarctic NE Pacific have emphasized the importance of iron as a limiting micronutrient of primary production, particularly by large diatoms (Martin and Gordon 1988; Martin and Fitzwater 1988; Martin et al. 1989; Coale 1991; Boyd et al. 1996). A biochemical approach has also demonstrated that diatoms are iron-stressed in the NE Pacific (La Roche et al. 1996). These findings have been very valuable in advancing the hypothesis of iron-limitation hypothesis.

On the other hand, the iron-limitation status is not simple in the subarctic NW Pacific. The western subarctic phytoplankton communities are usually moderate in biomass (0.5–1 μg Chl a L–1), but blooming of microphytoplankton, such as diatoms, sometimes occur during spring (Odate and Maita 1988/1989). Although the annual input of atmospheric iron in the NW Pacific is higher than that in the NE Pacific, most of the dust storms in Asia occur during spring, and atmospheric iron input could be low during summer and autumn (Tsunogai et al. 1985; Donaghay et al. 1991). Dissolved iron concentrations of <0.2 nM have been reported in May in the Oyashio region (Nishioka et al. 2003), and the HNLC condition appears after the termination of spring bloom in this region (Saito et al. 2002). These observations suggest that iron-limitation status in the NW Pacific varies spatiotemporally.

In the subarctic NE Pacific, the iron-limitation status also changes between spring/summer and winter. Maldonado et al. (1999) reported that growth and photosynthesis of phytoplankton are enhanced in winter by increasing either irradiance or iron, suggesting co-limitation of algal community by both the resources. In addition to the decrease in light availability during winter, temperatures in the surface mixed layer decrease from 14°C in summer to 3°C in winter. The growth rate of phytoplankton is a function of temperature (Eppley 1972), and elemental composition of algal cell is also influenced by temperature (Goldman 1979). Thus, iron nutrition, light, and temperature could have combined effects on the physiology and growth of phytoplankton.

In this study, the effect of iron on the growth of phytoplankton community was examined during spring and autumn in the NW Pacific, as well as under the different light and temperature conditions during winter in the NE Pacific, to reveal the differences in the community response to iron under various environmental conditions.

4-1. Experimental conditions of onboard iron-enrichment bottle incubation experiments

In the subarctic NW Pacific, iron-enrichment experiments were conducted with indigenous plankton communities at an oceanic region (Station A, 45°N; 165°E), during R/V Hakuho-Maru KH-93-4 cruise in October 1993, and at the Oyashio region (Stations A4, A7, A17) during R/V Wakataka-Maru WK0306 cruise in May 2003. Trace-metal clean technique was used as described by Takeda and Obata (1995).

Water samples with their resident phytoplankton were collected from the surface mixed layer at a depth of 15 m (∼40% I0) at Station A using acid-cleaned polypropylene bellows pump with silicone tubing. An acid-cleaned 20-L polyethylene carboy was filled with water to homogenize the water. The measured initial concentrations of dissolved iron in the samples were 0.22 nM at Station A. The sample water was dispensed into replicate 1-L polycarbonate incubation bottles and enriched with 1 nM FeCl3. The control bottles received no addition. The water samples were not prescreened with netting to exclude grazing organisms from incubation bottles. The bottles were incubated on deck in running surface seawater tanks to maintain the surface seawater temperatures (10–13°C at Station A) for 5 days. The incubation baths were covered with neutral density screens, which provided shading to 40% of the sea surface light level. During the course of the incubations, two bottles were withdrawn for each treatment from the incubation tanks at a time and chlorophyll-a (size fractionation at 3 and 10 μm) nutrients, and dissolved iron were measured. Initial samples were collected directly from 20-L carboys. Net specific growth rate was calculated by a linear regression of the chlorophyll-a increase during Days 1, 3, and 5.

In the Oyashio region of the subarctic NW Pacific, the surface water samples were collected from 10-m depth using X-Niskin sampler suspended on a Kevlar hydro-wire at Station A4 (42°15'N; 145°08'E), A7 (41°30'N; 145°30'E), and A17 (39°N; 146°45'E) during May 7–19, 2003. Sample waters were spiked with 4 nM FeCl3 or 300 nM DFB. Control bottles received no addition of iron or DFB. The fungal siderophore DFB was used to suppress iron uptake by phytoplankton in the incubation bottles, which allowed examining the existence of biologically available iron by comparing the growth with unamended control. The treated samples and controls were incubated for 5 days using triplicate 4-L polycarbonate carboys (A4) or duplicate 1-Lpolycarbonate bottles (A7 and A17) in temperature-controlled water tanks covered with neutral density screens, which provide shading to 30% of the incident light. The water samples were prescreened with 202-μm Teflon mesh netting to exclude grazing organisms from the incubation bottles.

In the subarctic NE Pacific, bottle incubation experiments were performed at Ocean Station P (50°N; 145°W) during CCGS J. P. Tully cruise 9803 in February 1998. Water samples were collected from 40-m depth using Teflon pump with Teflon tubing. Very low (0.07 nM) ambient concentration of total dissolvable iron was measured for the sample water. Six treatments were compared (Table 6). Two different light and temperature conditions were selected to mimic in situ condition in summer and winter. The sample seawater in 500-ml polycarbonate bottles was placed in temperature-controlled water baths illuminated by white fluorescent light tubes and sacrificed for analyses after 4 and 7 days of incubation periods. Duplicate bottles were used for each treatment. The water samples were not prescreened with netting to exclude grazing organisms from incubation bottles. Concentrations of chlorophyll-a and macronutrients as well as the composition of the phytoplankton species were analyzed for each bottle as well as the initial samples.


Table 6. Experimental conditions for light, temperature, and iron treatments and control during the bottle incubation experiments at Ocean Station P (50°N, 145°W) in February 1998.

T+L — high-temperature and high-light; T+L+Fe — high-temperature and high-light with iron addition.

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4-2. Effects of iron enrichment on phytoplankton assemblage in the Western Subarctic Gyre

At the subarctic Station A, nitrate concentrations in excess of 10 μM were observed in the surface mixed layer, where chlorophyll-a concentrations ranged from 0.9 to 1.3 μg L–1. The primary productivity measured at Station A during the same cruise (864 mgC m–2d–1, Odate and Furuya 1995) is similar to that reported in the subarctic NE Pacific (727 mgC m–2d–1, Frost 1993). Vertical profile of total dissolvable iron showed nutrient-like distributions with surface depletion at <0.05–0.2 nM. This pattern is similar to the previous observation in the Gulf of Alaska (Martin et al. 1989). Iron levels remained constant in deep waters below 1000 m, but were considerably high (>1.3 nM), when compared with the average concentration of 0.76 nM in the North Pacific and other open oceans (Johnson et al. 1997).

Addition of 1 nM iron to the surface water with its resident phytoplankton induced chlorophyll-a increase relative to the controls after 1 day (Fig. 17). The final chlorophyll-a concentrations, expressed as percent of the initial values, were 425% for the iron-enriched sample and 138% for the control. Although the 0.7–3-μm-size fraction initially formed the largest proportion (45%) of the total chlorophyll-a biomass, a remarkable increase in the >10– and 3–10-μm-size fractions was observed in the iron-enriched samples. After 5 days of incubation, the >10-μm-size fractions comprised 48% of the total chlorophyll-a biomass. In fact, 74% of the chlorophyll-a increase in the iron treatment was due to the increases in the >10– and 3–10-μm phytoplankton. These results were, in general, similar to those observed in the previous enrichment experiments performed in the Gulf of Alaska (Martin and Fitzwater 1988; Martin et al. 1989; Coale 1991; Boyd et al. 1996), and both indicate iron limitation of phytoplankton productivity in the subarctic North Pacific.


Fig. 17. Time-course change in chlorophyll a concentrations in three size fractions in controls and Fe-enriched bottles during on-deck incubation at Station A in the subarctic NW Pacific. Dissolved iron concentration in the control bottles was 0.22 nM. Values are the mean of duplicate samples.

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Microscopic observations of the iron-enriched samples showed that large diatoms dominated the biomass at the end of the incubation. Centric diatoms, such as Thalassiosira spp. (cell diameter of about 30 μm) accounted for most of the phytoplankton biomass in the >10-μm-size fraction. In the 3–10-μm-size fraction, the important taxa were pennate diatoms (cell size of about 5 × 40 μm) and coccolithophores (cell diameter of about 8 μm).

The net phytoplankton growth rates in the incubation bottles were significantly higher in bottles enriched with iron (Table 7). Although the net growth rate of the 0.7–3-μm-size fraction was not as high as that of the >10– and 3–10-μm-size fractions, the iron- mediated increases in the net growth rates were significant in all the size fractions (t-test, P < 0.05). The observed rates of 0.26–0.43 d–1 in iron-enriched samples are similar to those observed at the Gulf of Alaska (Martin et al. 1991) and during the SERIES experiment (Marchetti et al. 2006c), while diatom species appear to grow at faster rates in other experiments (Coale 1991; Boyd et al. 1996). The slower rate of net increase in 0.7–3-μm phytoplankton may reflect the fact that microzooplankton grazing regulates the biomass of these small cells (Frost 1991; Miller et al. 1991), as the growth rates of ciliates were found to be high (1.5–1.6 division d–1), and their food requirement was observed to correspond to 56% of the total primary production in this station (Kato 1995).


Table 7. Net growth rates of phytoplankton biomass during the iron enrichment experiments in the subarctic NW Pacific (Station A).

*The rate difference between the control and iron treatment is statistically significant (P < 0.01).

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Increased utilization of nitrate and phosphate was observed in the iron-enriched samples than in the controls. However, the net decrease in these nutrients with iron addition was only 2.3–2.6 μM nitrate and 0.17–0.27 μM phosphate during 5 days of incubation. In the case of silicic acid, about 3 μM was consumed both in the iron-enriched samples and the controls. As a result, the molar consumption ratios of silicic acid to nitrate and silicic acid to phosphate in the control bottles (2.6Si:1N and 31Si:1P) were twice as high as those in the iron-enriched bottles (1.2Si:1N and 14Si:1P). These results suggest that resident diatom populations were iron-stressed in October (Takeda 1998), although changes in diatom composition toward species that are inherently more lightly silicified could partly contribute to the observed changes in Si:N and Si:P ratios after iron enrichment (Marchetti and Cassar 2009). Concentrations of ammonia also decreased from 0.54 to <0.1 μM during the first 3 days of incubation, both in the iron treatments and the controls.

The total dissolvable iron concentration in the iron-enriched samples decreased linearly during the entire incubation, suggesting that removal of iron from the dissolved phase took place at nearly constant rate during the experiment, in spite of the remarkable increase in algal biomass after Day 3 (Fig. 18). The decrease in the total dissolvable iron concentration during the 5 days of incubation was found to be 54% of the added iron (Table 8). However, in the controls, no significant change in the total dissolvable iron concentration was observed. The proportions of added iron adsorbed on the bottle walls were determined by rinsing the inner wall of the incubation bottles with formic acid-ammonium formate buffer solution (pH 3) at the end of the experiments, and it was only 1.6–2.6% even after the pronounced increase in biomass in iron-enriched bottles.


Fig. 18. Changes in dissolved iron concentration in incubation bottles enriched with iron (●) and in controls (○) during on-deck incubation at Station A in the subarctic NW Pacific. The error bars are the standard deviation of the duplicate samples.

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Table 8. Dissolved iron concentration at the initial time point (Day 0) and final time point (Day 5) in the controls and iron-enriched samples during the iron enrichment experiments in the subarctic NW Pacific (Station A). Decrease percentage is calculated as 100 × (initial Fe – final Fe)/initial Fe. The ratios of Fe decrease to particulate organic carbon increase (Fe:POC) in the control and iron-enriched samples are also shown. The values are the mean of duplicates and the standard deviation.

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Addition of iron produced twice as much POC as that of the controls during 5 days of incubation (Table 9). The ratio of POC:Chlorophyll-a at the end of the incubation was lower in the iron-enriched samples, indicating that iron stimulated chlorophyll-a synthesis more strongly than carbon fixation in the subarctic samples. As a large portion of extracellular iron could be detected while measuring the total dissolvable iron at pH 3.2 (Nishioka and Takeda 2000), the observed decrease in the total dissolvable iron was assumed to be equivalent to the intracellular iron taken up by the phytoplankton. The ratio of total dissolvable iron consumed to the total community POC produced (Fe:C) in the iron-enriched bottles was about 4.5 times higher than that observed in the controls (Table 8). The observed Fe:C ratio (54 μmol Fe mol C–1) was 5 times higher than the cellular Fe:C ratios needed for maximum growth of oceanic phytoplankton (Sunda et al. 1991; Sunda and Huntsman 1995), suggesting that phytoplankton cells in the iron-enriched bottles had elevated iron quota, perhaps by luxury iron uptake and storage. Marchetti et al. (2006a) reported that oceanic centric diatom T. oceanica could store about 20 times higher amounts of intracellular iron under high iron concentrations, although oceanic Pseudo-nitzschia species have demonstrated much larger capacity for iron storage by using ferritin (Marchetti et al. 2009). Thus, the slower rates of iron removal per unit of chlorophyll observed in the latter phase of incubation might result from a decrease in the iron uptake rate of the phytoplankton assemblage.


Table 9. Initial (Day 0) and final (Day 5) concentrations of POC and Chl a and the ratio of POC/Chl a for the controls and iron-enriched samples during the iron enrichment experiments in the subarctic NW Pacific (Station A). The values are the means of duplicate samples and the standard deviation.

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In the subarctic NE Pacific, Boyd et al. (1996) reported that >18-μm-size fraction accounted for >80% of algal stocks by Day 6 in the iron-enriched carboys, whereas in the present study, cells of >10 μm in size were responsible for 48% of chlorophyll-a biomass by Day 5 in the iron treatment at Station A. On the other hand, the effect of iron enrichment on the growth of small (0.7–3 μm) phytoplankton was evident in the present study, but small changes in biomass and growth rates were observed for the small (<5 μm) phytoplankton cells during the iron-enrichment experiments carried out by Boyd et al. (1996). These differences in the relative importance of large cells may be partly due to the different seed populations, as iron enrichment mainly stimulated the growth of large centric diatoms (Thalassiosira sp.) in the subarctic NW Pacific and pennate diatoms in the subarctic NE Pacific (Martin and Fitzwater 1988; Coale 1991; Boyd et al. 1996; Tsuda et al. 2005; Marchetti et al. 2006c). Calculations of the diffusional flux and iron uptake kinetics at the cellular level revealed that pennate diatoms having long and narrow cell shape are better competitors than rounded centric diatoms under low iron conditions (Wells 2003; Marchetti and Cassar 2009).

4-3. Effects of iron enrichment on phytoplankton assemblage in the Oyashio region

In the spring of 2003, phytoplankton blooming of the Oyashio region started in late April. Thus, the bottle incubation experiments in May were conducted under several different nutrient conditions and phytoplankton stocks (Table 10). The northern Station A4 is close to coast and most influenced by Oyashio current, while the southern Station A17 is affected by subtropical water to some extent (Fig. 8). Initial chlorophyll-a concentrations were high at 4.1–16 μg L–1, reflecting the developing phase of a spring bloom at Station A7 and bloom decline phase at Stations A4 and A17. Phytoplankton assemblages were dominated by large (>10 μm) cells and mainly comprised diatoms and cryptophytes. Dissolved iron concentrations were around 0.1 nM at Stations A7 and A17, but a high value was observed at Station A4.


Table 10. Initial conditions at Stations A4, A7, and A17 in the Oyashio region in May 2003. Reprinted from PhD Thesis, Yoshiko Kondo, Dynamics of the organic Fe complexing ligands and phytoplankton in the Pacific Ocean, 256 pp., © 2007.

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At Station A4, there was only 1.2 μM nitrate at the beginning of the experiment, and it was used up within 1 day in all treated and untreated bottles (Fig. 19). Due to this macronutrient limitation, chlorophyll-a concentration remained at a constant level or gradually decreased during 5 days of incubation (Fig. 20). The DFB treatment did not have an apparent effect on changes in chlorophyll-a, nitrate, and phosphate but decreased the consumption of silicic acid in the bottles, indicating some restriction of diatom growth by organic complexation of the available iron.


Fig. 19. Time course of nitrate (a), phosphate (b), and silicic acid (c) concentrations in controls (●) and incubation bottles enriched with iron (□) or DFB (△) during on-deck incubation at Station A4 in May 2003. The values are the means of triplicate samples, and the error bars are the standard deviation. Reprinted from PhD Thesis, Yoshiko Kondo, Dynamics of the organic Fe complexing ligands and phytoplankton in the Pacific Ocean, 256 pp., © 2007.

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Fig. 20. Changes in chlorophyll a concentration in two size fractions in controls (a) and incubation bottles enriched with iron (b) or DFB (c) during on-deck incubation at Station A4 in May 2003. The values are the means of triplicate samples, and the error bars are the standard deviation. Reprinted from PhD Thesis, Yoshiko Kondo, Dynamics of the organic Fe complexing ligands and phytoplankton in the Pacific Ocean, 256 pp., © 2007.

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At Station A7, iron enrichment clearly enhanced the growth of both small- and large-size fractions of phytoplankton (Fig. 21). Control bottles also showed some increase in chlorophyll-a concentration during 5 days, while changes in chlorophyll-a and macronutrients were minimum after DFB treatment (Fig. 21, 22). These results suggest that there was limited amount of biologically available iron to support gradual growth of phytoplankton, although the measured initial dissolved iron concentration was only 0.06 nM. Consumption ratio of silicic acid to nitrate ([Si(OH)4]:[NO3]) in the control (1.7) was between that obtained after iron treatment (1.2) and DFB treatment (2.7). As the [Si(OH)4]:[NO3] ratio of diatoms was found to increase up to 2–3 under iron limitation (Takeda 1998), the observed ratio also indicates that the resident phytoplankton population was stressed by the low iron content.


Fig. 21. Changes in chlorophyll a concentration in two size fractions in controls (a, d) and incubation bottles enriched with iron (b, e) or DFB (c, f) during on-deck incubation at stations A7 (a–c) and A17 (d–f) in May 2003. The values are the means of triplicate samples. Reprinted from PhD Thesis, Yoshiko Kondo, Dynamics of the organic Fe complexing ligands and phytoplankton in the Pacific Ocean, 256 pp., © 2007.

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Fig. 22. Time course of nitrate (a, d), phosphate (b, e), and silicic acid (c, f) concentrations in controls (●) and incubation bottles enriched with iron (□) or DFB (▲) during on-deck incubation at Station A7 (a–c) and A17 (d–f) in May 2003. The values are the means of triplicate samples. Reprinted from PhD Thesis, Yoshiko Kondo, Dynamics of the organic Fe complexing ligands and phytoplankton in the Pacific Ocean, 256 pp., © 2007.

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At Station A17, small increases in chlorophyll-a concentration and nitrate consumption rate were observed in the iron treatment, but these changes were restrictive, owing to the very low initial nitrate concentration (Fig. 21, 22). Relatively high chlorophyll-a concentration was observed 2 days after DFB treatment, but it decreased to a low level similar to the control after 5 days. The reason for this fluctuation is not clear.

These bottle incubation experiments at three different conditions highlight the complex future of iron-limitation status in the Oyashio region during spring. Phytoplankton population at the northern coastal Station A4 was found to be controlled by the availability of macronutrients, mainly nitrate, while both nitrate and iron were observed to limit the phytoplankton growth at the southern Station A17. At Station A7, blooming of resident phytoplankton was stressed by iron deficiency, and diatoms slowly consumed macronutrients with high Si:N ratio, which caused depletion of silicic acid before the consumption of nitrate. The situation observed at Station A7 could explain the appearance of HNLC waters after the termination of spring bloom (Saito et al. 2002); however, both the quantity and composition of the remaining macronutrients could be altered depending on the iron availability. These results suggest that the balance between macronutrients and iron supply principally controls the magnitude and persistence of spring bloom in the Oyashio region.

4-4. Limitation of phytoplankton growth by iron, light, and temperature in the subarctic NE Pacific during winter

Bottle incubation experiments were performed using ambient surface water in February 1998 at Ocean Station P in the subarctic NE Pacific. Other iron- enrichment experiments conducted on the same cruise along the Line P at Station P4 (48°39'N; 126°40'W), P16 (49°17'N; 134°40'W), and P26 (Ocean Station P) showed that the growth of ambient phytoplankton at Ocean Station P and Station P16 was enhanced by 1 nM iron addition, but there was no effect of iron- enrichment at the coastal Station P4 (data not shown). Although these experiments did not mimic the low light condition by the deep surface mixing layer, the observed iron stress at the oceanic stations is consistent with the gradient of dissolved iron concentration in the surface water from 1.7 nM at Station P4 to 0.01–0.12 nM at Station P16 and Ocean Station P (W. K. Johnson, unpublished data).

Figure 23 shows the changes in size-fractionated chlorophyll-a concentrations during the experiments. The controls exhibited little growth under low-light (42 μmol m–2s–1; 8:16 h LD cycles), low-temperature (6°C), and low-iron concentration (0.07 nM). Iron addition (1 nM) or high-temperature (13°C) treatment alone did not show clear increase in chlorophyll-a concentration, when compared with the controls, while chlorophyll-a production was up by 1.5 folds under the high-light condition (170 μmol m–2s–1; 14:10 h LD cycles). Combination of the high-light and high-temperature (T+L) showed a 2.3-fold increase in chlorophyll-a level. In these high-light conditions, chlorophyll-a concentration in 0.7–5-μm-size fraction increased significantly over the first 4 days of incubation and then steadily declined, while that in >5-μm-size fraction increased continuously. The high-temperature and high-light treatment associated with the addition of iron (T+L+Fe) showed the most pronounced effects on chlorophyll-a biomass both in 0.7–5- and >5-μm-size fractions (15.5-fold increase).


Fig. 23. Changes in chlorophyll a concentration in two size fractions in controls (low temperature, low light and low-iron) and incubation bottles treated with high temperature, high light, high iron, and combinations of those (T+L, high temperature and high light; T+L+Fe, high temperature and high light with iron addition). The experiments were conducted onboard using surface water collected at Ocean Station P in February 1998. See Table 6 for the detailed experimental conditions.

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The sample seawater in the incubation bottles initially contained high concentration of nutrients: 13 μM nitrate, 1.2 μM phosphate, and 20 μM silicic acid. Phytoplankton in the T+L+Fe treatment consumed about 5 μM nitrate, 0.3 μM phosphate, and 4 μM silicic acid during 7 days of incubation, but other treatments and control showed only a few decrease (<1 μM nitrate) due to the relatively low chlorophyll-a levels achieved in the bottles.

Small autotrophic cells and Synechococcus spp. were the dominant groups in the water sample before the incubation (Table 11). Some pennate diatoms and small prymnesiophytes were also present, but only a few centric diatoms were observed in the initial samples. High-temperature treatment increased the abundance of initially dominant small autotrophs and Synechococcus, while pennate diatoms became the major group in the high-light treatment. The population of Synechococcus showed weak inhibition in the high-light treatment. Iron addition had no apparent effect on phytoplankton abundance, when compared with the control. The combination of high temperature and high light (T+L treatment) increased the cell numbers of all major phytoplankton groups, but the largest increase was observed in the T+L+Fe treatment. In the >5-μm-size class, pennate diatoms were the most dominant group in the T+L and T+L+Fe treatments, while centric diatoms and photosynthetic dinoflagellates also contributed to some extent.


Table 11. Initial and final cell densities (cells ml–1) for each size class of phytoplankton and microzooplankton during the bottle incubation experiments at Ocean Station P (50°N, 145°W) in February 1998.

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Microscopic observation of the plankton samples also indicated significant increase in microzooplankton in the T+L and T+L+Fe treatments (Table 11). High microzooplankton grazing pressure seems to be responsible for the observed decrease in 0.7–5-μm-size chlorophyll-a in the T+L treatment after 4 days, while small phytoplankton in the T+L+Fe treatment could overcome the microzooplankton grazing owing to their growth stimulation by the addition of iron.

The observed results give strong evidence that growth of subarctic NE Pacific phytoplankton was co-limited by light, temperature, and iron in winter, although iron alone could not have pronounced effect within 7 days of incubation. Within these three resources, light seems to be an important limiting factor for algal growth, especially in pennate diatoms. As most of the iron is involved in photosynthesis, an increase in irradiance reduces the cellular iron requirement of photoautotrophs (Sunda and Huntsman 1997; Maldonado et al. 1999). This reduction in iron requirement could partly explain the observed increase in the growth of pennate diatoms in the high-light treatment. However, iron addition alone did not stimulate the growth of pennate diatoms and other phytoplankton, although a little increase in chlorophyll-a concentration without distinguishable change in the cell number suggests some enhancement of pigments synthesis by added iron. Therefore, light should be considered as a primary factor for algal biomass accumulation in winter.

In spring and summer, with the increase in temperature and light availability in the surface mixed layer, both small and large phytoplankton may increase their stock to some extent. However, microzooplankton respond rapidly to the increase in small phytoplankton abundance, and the grazing balances the growth of small cells. The observed experimental results indicate that additional iron is needed for small phytoplankton growth to overcome the microzooplankton pressure under the high-temperature and high-light condition. Therefore, iron is an important factor for both small and large phytoplankton to increase their stocks at blooming levels and to use up the macronutrients existing in the surface water during summer.

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5. Biogeochemical responses of planktonic ecosystems during three mesoscale iron-enrichment experiments in the subarctic North Pacific

To test the iron limitation hypothesis on the scale of the entire community, three mesoscale iron-enrichment experiments were conducted in the subarctic North Pacific Ocean during summer. The first experiment in the Western Subarctic Gyre was the SEEDS in 2001 (Tsuda et al. 2003), followed by the SERIES performed near Ocean Station P in the Alaska Gyre in 2002 (Boyd et al. 2004). The second experiment in the Western Subarctic Gyre, SEEDS-II, was carried out at almost the same location as SEEDS in 2004 (Tsuda et al. 2007). All the three enrichment experiments have confirmed that iron availability strongly influences primary productivity and food-web structure. However, SEEDS differed from the other two experiments by its massive bloom of a chain-forming centric diatom. SERIES, on the other hand, caused an initial bloom of small phytoplankton followed by a larger bloom of various diatom species and provided unique information on the fate of carbon after the decline of the iron-induced phytoplankton bloom. SEEDS-II resulted in a much smaller buildup of phytoplankton biomass, in which small non-diatom species dominated, probably due to a higher initial mesozooplankton biomass and grazing pressure on the diatoms.

In this section, a summary of key results obtained in these experiments has been presented, which highlights the major differences observed in the biological and geochemical responses and discusses the reasons for the unexpected response observed during the last experiment, SEEDS-II. This intercomparison has been made to understand the effect of iron on the biogeochemical cycles and the complexity of ecosystem responses to iron in the eastern and western subarctic HNLC waters.

5-1. Summary of biological and geochemical findings in SEEDS, SERIES, and SEEDS-II

The first mesoscale in situ iron-enrichment experiment in the subarctic North Pacific, SEEDS, was performed by fertilizing a patch of 8 × 10 km with iron at 48°30'N, 165°E near the center of the Western Subarctic Gyre on July 18, 2001 (Day 0 of the experiment), and the observation continued until August 1, 2001 (Fig. 24). In the eastern subarctic North Pacific, SERIES was started on July 9, 2002 (Day 0) by releasing iron to an area of 8.5 × 8.5 km at 50°12'N, 144°27'W. A second release of iron was carried out in response to declining dissolved iron levels on July 16, 2002 (Day 7), and the patch was followed until August 4, 2002 (Day 26) by a three-ship operation. The second experiment in the western subarctic North Pacific, SEEDS-II, took place from July 20 (Day 0) to August 15, 2004 (Day 26) at 48°N, 166°E, 93-km southeast of the SEEDS site, and observed by two research vessels. Iron was released twice to an area of 8 × 8 km on Day 0 and to the patch on July 20, 2004 (Day 6).


Fig. 24. Location of the experimental sites for SEEDS, SEEDS-II in the Western Subarctic Gyre, and SERIES near the Ocean Station P (P26) in the Alaska Gyre.

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These experiments were conducted in HNLC waters under similar subarctic summer conditions, but there were some differences in the physical, chemical, and biological characteristic of the waters (Table 12). SEEDS had shallow surface mixed layer of <10 m. The initial phytoplankton population in SERIES was dominated by small algal groups, and the biomass was lower, when compared with SEEDS and SEEDS-II.


Table 12. Comparison of the initial conditions in the surface water among SEEDS, SERIES, and SEEDS-II.

*Fv/Fm: photochemical quantum efficiencies of algal photosystem II. Data are from Takeda and Tsuda (2005), Boyd et al. (2005), and Tsuda et al. (2007).

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Similar amounts of iron as FeSO4 dissolved in acidified seawater were released to the surface water in these experiments, but theoretical increase in iron concentration was low in SEEDS-II because of the deeper mixed layer depth when iron was released (Table 13). In SEEDS, the actual measured concentration of dissolved iron at the center of the iron-enriched patch was nearly 3 nM on Day 1 and the concentration decreased gradually and continuously until the end of the observation (Fig. 25a). In SERIES, dissolved iron increased up to 2.2 nM and showed decreasing pattern similar to that of SEEDS. However, the initial increase in dissolved iron during SEEDS-II was only half of that observed in SEEDS, and the concentration decreased quickly. Although the dissolved iron concentration was recovered by the second iron release, it decreased again to the initial background level after 2 weeks.


Table 13. Times and quantities comprising the iron infusion in SEEDS, SERIES, and SEEDS-II experiments.

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Fig. 25. A comparison of changes in surface dissolved iron (a) and chlorophyll a (b) concentrations in the surface water during the time course of SEEDS, SERIES, and SEEDS-II. The data are from Tsuda et al. (2003), Boyd et al. (2004), Wong et al. (2006b), Tsuda et al. (2007), and Nishioka et al. (2009).

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Each experiment demonstrated that iron supply could improve the physiological state of ambient phytoplankton in a short duration. Photochemical quantum efficiencies of algal photosystem II (Fv/Fm) were observed to increase within 24 hours and showed a peak value of 0.44 on Day 9 in SERIES (Boyd et al. 2004). However, in the experiments in the western subarctic Pacific, it took 2 days (SEEDS-II) or 3 days (SEEDS) for phytoplankton to show physiological response to the addition of iron. The maximum value of Fv/Fm was observed on Day 9 in SEEDS (0.44) and on Day 13 in SEEDS-II (0.42) (Tsuda et al. 2003; Suzuki et al. 2009).

Increase in chlorophyll-a concentration became apparent in the surface water after Day 4 in SEEDS, and chlorophyll-a accumulated rapidly in the surface water to very high levels more than 15 μg L–1 (Fig. 25b). In SERIES, chlorophyll-a showed gradual increase, reached about 5 μg L–1 on Days 16–18, and then decreased. Chlorophyll-a concentration in SEEDS-II increased from Day 4, peaked at 3 μg L–1 between Days 11 and 13, and returned to the initial level on Day 20. Considering the difference in the surface mixed layer depth, comparison of water column integrated chlorophyll-a concentrations showed that the maximum value of SEEDS (240 mg m–2) and SERIES (130 mg m–2) are 3.3 and 1.8 times higher than that of SEEDS-II (74 mg m–2), respectively (Tsuda et al. 2007). The maximum primary productivity in the surface mixed layer was 310 mg C m–3d–1 in SEEDS (Kudo et al. 2005), 180 mg C m–3d–1 in SERIES (Marchetti et al. 2006c), and 112 mg C m–3d–1 in SEEDS-II (Kudo et al. 2009).

The magnitude of nitrate consumption showed a similar aspect to that of chlorophyll-a (Fig. 26). However, a time-series comparison of nitrate and chlorophyll-a data indicated that a large portion of nitrate decrease occurred without increase in chlorophyll-a after Day 5 during SEEDS-II. Such nitrate consumption out of proportion to chlorophyll-a increase was only observed during the peak of bloom in SEEDS and SERIES. Ammonium uptake dominated phytoplankton's nitrogenous nutrition by Day 17 in SEEDS-II (Kudo et al. 2009). Silicic acid almost depleted when chlorophyll-a reached the maximum in SERIES, but a decrease in silicic acid was only about 4 μM in SEEDS-II (Fig. 26). The drawdown ratio of silicic acid to nitrate doubled when chlorophyll-a reached near the maximum value both in SEEDS and SERIES (Boyd et al. 2005; Kudo et al. 2005). This suggests that at the later phase of bloom development, diatoms were under physiological stress from scarce iron (Takeda 1998), or there were morphological changes in the frustules of the dominant pennate diatoms in response to iron availability during SERIES (Marchetti and Harrison 2007; Marchetti and Cassar 2009). In SEEDS, decrease in light availability due to self-shading effect could also have had an effect on the drawdown ratio of silicic acid to nitrate (Davis 1976; Saito and Tsuda 2003).


Fig. 26. A comparison of changes in surface concentrations of nitrate (a) and silicic acid (b) during the time course of SEEDS, SERIES, and SEEDS-II. The data are from Tsuda et al. (2003), Boyd et al. (2004), and Tsuda et al. (2007).

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The development of the bloom accompanied floristic shifts in phytoplankton communities. In SEEDS, the initial assemblage having high abundance of small picoeukaryotes was replaced by a large chain-forming centric diatom, although almost all diatom species responded to the addition of iron (Tsuda et al. 2005). In SERIES, the biomass of pico- and nanophytoplankton increased first, followed by the large diatoms composed of several pennate and centric groups (Marchetti et al. 2006b). On the other hand, diatoms did not bloom in SEEDS-II, and autotrophic nanoflagellates, such as cryptophytes and prasinophytes, became predominant in the phytoplankton community (Suzuki et al. 2009). As a result, contribution of diatoms to chlorophyll-a biomass increased from about 30% to 70% in SEEDS and SERIES but gradually decreased during SEEDS-II (Fig. 27).


Fig. 27. A comparison of changes in contribution of diatoms to total chlorophyll a biomass during the time course of SEEDS, SERIES, and SEEDS-II. The data are from Suzuki et al. (2005), Wong and Crawford (2006), and Suzuki et al. (2009).

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During SEEDS, iron enrichment resulted in large drawdown of CO2 fugacity in surface water along with the accumulation of enormous algal biomass (Table 14). Around 78% of the fixed algal carbon remained in the mixed layer at the end of the observation on Day 13 of the SEEDS bloom (Tsuda et al. 2003). Indeed, the monitoring period during SEEDS was insufficient to determine the fate of the bloom. In the case of SERIES and SEEDS-II, the 26-day-long observations allowed the examination of the decline phases of the iron-induced phytoplankton blooms as well as the fate of organic carbon production, although the observed decreases in CO2 fugacity were smaller than those observed during SEEDS.


Table 14. A comparison of the fugacity of carbon dioxide (fCO2) in the unfertilized ambient water and the minimum value observed during the SEEDS, SERIES, and SEEDS-II experiments. Data are obtained from Tsuda et al. (2003, SEEDS), Wong et al. (2006a, SERIES), and Tsuda et al. (2007, SEEDS-II).

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Downward fluxes of POC intercepted at 40–50-m depth by sediment traps located near the center of the iron-fertilized patch are shown in Fig. 28. In SERIES, export flux of POC from iron-stimulated phytoplankton blooms increased to 40 mmol C m–2d–1 between Days 24 and 26. Although the accumulated mixed-layer POC declined by 79% during the observation, only 18–22% of this POC was exported to a depth of 50 m, while mesozooplankton herbivory accounted for 10% of the POC decline (Boyd et al. 2004). Majority of the iron-elevated mixed-layer POC was remineralized by bacterial activities both within and below the mixed layer. Similar high fluxes of more than 40 mmol C m–2d–1 were observed during the decline phase of the SEEDS-II bloom. However, Tsuda et al. (2007) pointed out that the POC flux was largely controlled by the copepod biomass, both inside and outside of the iron-fertilized patch, and that the iron had a minor effect on the export flux through growth stimulation of phytoplankton in SEEDS-II. About 33–43% of the primary production was exported as sinking particles, and 18–25% was transferred to mesozooplankton net growth (Tsuda et al. 2007; Kudo et al. 2009). The increase in POC export flux in the iron-fertilized patch was not significant, when compared with that outside of the patch in SEEDS. The export flux between Day 4 and Day 13 was 11% of the integrated primary production in the patch (Tsuda et al. 2003).


Fig. 28. A comparison of time-series variations of POC-sinking flux collected at 40- or 50-m sediment traps inside and outside of the iron-fertilized patch during SEEDS, SERIES, and SEEDS-II. The data are from Tsuda et al. (2003), Boyd et al. (2004), and Aramaki et al. (2009).

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These three mesoscale experiments have revealed a remarkable similarity with respect to the algal physiological properties and the general bloom trends, and the findings are very useful in providing a comprehensive view of the role of iron as a limiting nutrient for phytoplankton growth in the subarctic North Pacific. However, there are several significant departures in biological signals and floristic shifts, which may disclose weaknesses and gaps in the understanding of the functioning of this ecosystem.

5-2. Comparison of SEEDS and SEEDS-II: Why diatoms could not dominate in the iron-induced SEEDS-II bloom?

To explain the observed contrasting responses between SEEDS and SEEDS-II, high mesozooplankton grazing in SEEDS-II was proposed by Tsuda et al. (2007). The mesozooplankton biomass at the beginning of the experiments was 2.8 and 2.6 times higher than that of SEEDS and SERIES, respectively (Tsuda et al. 2007). Although the reason for the higher initial mesozooplankton biomass in SEEDS-II is unknown, the mesozooplankton, dominated by a copepod Neocalanus plumchrus, increased their biomass exponentially by growth supported by selective feeding on microparticles, including diatoms (Tsuda et al. 2007, 2009; Saito et al. 2009). This mesozooplankton ingestion was accounted for 45% of the whole primary production (Kudo et al. 2009). The copepod biomass in the iron-fertilized patch reached the maximum 4 days earlier to a level 1.6 times greater than that observed outside the patch, and then suddenly decreased by the ontogenetic downward migration of N. plumchrus (Tsuda et al. 2007). In addition, the export flux of particulate silica in the iron-fertilized patch was 2.7 times higher than that of the outside of the patch (Saito et al. 2009), and the sinking particles collected at 40-m depth in the patch had ∼2.6 times higher Si:C ratio, when compared with the outside (Aramaki et al. 2009). On the other hand, a heterotrophic dinoflagellate, Gyrodinium sp., contributed to the loss of large diatoms during the later phase of SEEDS bloom (Saito et al. 2006a), but the dinoflagellate biomass decreased during SEEDS-II (Tsuda et al. 2009). These findings suggest that copepod grazing prevented the formation of an extensive diatom bloom and advanced the floristic shift to a picophytoplankton-dominated community. However, this is in contrast to the other iron-enrichment experiments where intense grazing on picophytoplankton prevented small cells from increasing in biomass and allowed diatoms to bloom. In SEEDS-II, the trophic cascading effects of the copepod grazing on nanoflagellates as well as microzooplankton have been proposed as a possible controlling mechanism for picophytoplankton abundance (Tsuda et al. 2009).

In addition to mesozooplankton grazing, other factors could also cause the relatively small response of diatoms in SEEDS-II, when compared with SEEDS. These include the difference in the initial seed populations, low supply of iron for luxury iron uptake by large diatoms, a deeper surface mixed layer depth, and iron limitation induced by the release of organic iron-complexing ligands by plankton assemblage.

As the existence of neritic centric diatom, Chaetoceros debilis, brought most of the increase in algal biomass in SEEDS, the absence of neritic fast-growing diatoms could have had some influence on the magnitude and timing of the bloom in SEEDS-II. However, in the phytoplankton assemblage of SEEDS-II site, there were oceanic diatoms similar to those that largely contributed to the development of the SERIES bloom (Tsuda et al. 2007). Therefore, difference in the initial seed populations cannot explain why large diatoms did not bloom during SEEDS-II. De Baar et al. (2005) proposed that the depth of surface mixed layer was important in determining the maximum chlorophyll-a level, maximum nitrate removal, and maximum fCO2 drawdown in the past in situ iron fertilization experiments. Again, the initial surface mixed layer depth in SEEDS-II is comparable with that in SERIES, and thus, the observed differences between these two experiments require reasons other than the surface mixed layer depth.

During SEEDS-II, the dissolved iron levels at 5-m depth were <0.3 nM throughout the experiment, except just after the first and second iron releases (Fig. 25a). The rapid disappearance of the injected iron from the surface water was caused by the strong dilution effect during the first 10 days of SEEDS-II, which was at least 5 times greater than that of SEEDS (Tsumune et al. 2009). In addition, three times deeper surface mixing at the beginning of SEEDS-II than that of SEEDS could also have led to low dissolved iron in the surface water (Table 12). In addition, nearly half of the iron in the surface mixed layer appeared to be lost between Days 9 and 11, when the surface mixed layer depth decreased to 15 m (Nishioka et al. 2009). With regard to the half-saturation constant for the growth of large algal cells obtained at Ocean Station P and during SEEDS (0.57–0.58 nM; Table 5), the low iron concentrations of <0.3 nM during SEEDS-II would not have been adequate to support substantial growth of large diatoms, although the dissolved iron concentrations in the patch exceeded the diffusion-limited threshold for rapid growth of pennate diatoms (Wells 2003). Concurrent observations of the increase in photosynthesis competence (Fv/Fm) of phytoplankton assemblage in the iron-fertilized patch and the continuous occurrence of flavodoxin (non-iron-containing flavoprotein) expressions for large-sized diatoms suggest that the iron enrichments restored algal cells from iron-deficiency stress, except for large diatoms (Suzuki et al. 2009). On the other hand, onboard incubation experiments conducted during SEEDS-II indicated that a certain amount of iron was available in the surface water until Day 11, or diatoms could grow to some degree by using the iron stored in the cell by luxury iron uptake, which might have occurred just after the iron release (Nishioka et al. 2009). Therefore, the observed difference in growth response between the phytoplankton assemblage in the iron-fertilized patch and that in the incubation bottles may be derived primarily from the absence of heavy grazing pressure on diatoms in the incubation bottles, where most of the mesozooplankton were removed by mesh screen at the water sampling. In addition, it appears that the iron infusions could not fulfill the requirements of large diatoms for their maximum growth, presumably due to the relatively low mixed-layer iron concentrations and/or reduction in its bioavailability by organic complexation.

In SEEDS-II, increases in iron-complexing organic ligands were observed after the first and second iron releases and at the decline phase of the bloom (Fig. 29). Kondo et al. (2008) presumed that the organic ligands were released by phytoplankton immediately after the first and second iron infusions, and zooplankton grazing on bacteria and phytoplankton might be another source of organic ligands for the second increase. The third increase may be explained by the siderophore production by iron-limited bacteria and phytoplankton cell lysis by virus infection, in addition to zooplankton grazing. Rapid increase in iron-complexing organic ligand concentration after the addition of iron had also been reported in other mesoscale experiments in the equatorial Pacific (Rue and Bruland 1997) and Southern Ocean (Boyé et al. 2005). Kondo et al. (2008) also interpreted horizontal dilution of iron-enriched patch water mass as a major process that caused decreases in the ligand concentrations after the first and second iron infusions. Thus, the released iron would have become complexed by these natural organic ligands. Existence of free organic ligands in the surface water may modify the biological availability of the iron not only in the in situ surface seawater but also in the onboard incubation bottles. As discussed in Section 3, bottle incubation experiments conducted on Day 17 indicated that the bioavailability of the added iron was reduced by up to ∼2 nM iron, and the existence of organic ligands as high as 2.1 nM in the surface water seems to be a likely reason for that (Fig. 16). In fact, Sato et al. (2007) showed that organic ligands formed during microzooplankton grazing reduced the iron bioavailability to phytoplankton and suppressed their growth. Other bottle incubation experiments conducted by Wells et al. (2009) during SEEDS-II indicated that diatoms had a poor ability to access iron in a hydroxamate-type siderophore, which is considered to be another source of organic ligands at the decline phase of the bloom. Although iron- limited diatoms may induce high-affinity iron uptake systems that can take up iron associated with siderophores (Maldonado and Price 1999; Maldonado et al. 2006), Wells et al. (2009) indicated that diatoms in the SEEDS-II patch lacked sufficient copper availability to activate these high-affinity systems. Therefore, complexation of iron by grazing-mediated ligands and/or siderophores appears to be one of the reasons for the prevention of dominance of large diatoms in the iron-fertilized patch water.


Fig. 29. Time-series measurements of the concentrations of iron-complexing organic ligands and dissolved iron in the surface water during SEEDS (a, Takeda and Kondo, unpublished) and SEEDS-II (b, Kondo et al. 2008). The error bars show the analytical errors.

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On the other hand, there was only a small increase in the concentration of iron-complexing organic ligands (<1 nM) after the iron release during SEEDS, and the measured concentrations were initially lower than and subsequently similar to the dissolved iron concentration in the surface water (Fig. 29), indicating that a part of the released iron would initially have been present as inorganic species and subsequently would have precipitated as a colloidal or particulate fraction in addition to the uptake by microorganisms. The difference in the concentrations and characteristics of organic ligands during SEEDS and SEEDS-II probably had a fundamental effect on the stability of iron in the dissolved fraction, as well as the biological availability of iron to the microbes (Hutchins et al. 1999b).

The findings here suggest that the growth rate of large diatoms was restricted to some extent by the decrease in the bioavailability of iron through its complexation with organic ligands released by plankton assemblage, as well as by the shortage of iron to fulfill the luxury iron uptake after iron infusions during SEEDS-II. A high initial mesozooplankton biomass and their feeding on microparticles, including diatoms, was a main factor that prevented the accumulation of diatom standing stocks in SEEDS-II. Both the deep surface mixed layer depth and absence of rapid-growing neritic diatoms could also contribute in determining the magnitude of the overall response. These are the factors that account for the large difference between the ecosystem responses observed in SEEDS and SEEDS-II. However, we still have limited information on the sources and sinks of natural organic ligands, and thus, their roles in the shift of dominant phytoplankton species during the iron-mediated bloom are the preferential subjects for future studies.

5-3. Changes in the nutrient drawdown ratio during iron-induced phytoplankton blooming

While there is a wide variation in Si:N ratios in diatoms independent of changes due to variable growth conditions, the average Si:N ratio of the diatoms has been reported to be about 1.0 (Brzezinski 1985). Several studies have observed a two- to three-fold increase in the Si:N and [Si(OH)4]:[NO3] uptake or consumption ratios in the iron-limited diatoms and diatom-dominated natural assemblages (e.g., Hutchins and Bruland 1998; Takeda 1998; Marchetti and Harrison 2007). As the nutrient supply ratio of [Si(OH)4]:[NO3] to the euphotic zone in the subarctic North Pacific ranges from 1.5 to 1.9 (Whitney and Freeland 1999; Saito et al. 2002), changes in the diatoms' [Si(OH)4]:[NO3] uptake ratio are important in determining the limiting macronutrient at the development of diatom-dominated blooms.

In SEEDS and SERIES, both nitrate and silicic acid were consumed with the development of iron-induced diatom blooms. During the later phases of these blooming, the [Si(OH)4]:[NO3] drawdown ratio in the surface mixed layer increased, and it directed the phytoplankton nutrition status towards the depletion of silicic acid (Fig. 30).


Fig. 30. Relationship between nitrate and silicic acid concentrations at 5- to 10-m depth inside and outside of the iron-enriched waters measured during SEEDS and SERIES. The lines and Si/N ratio designate the linear regression lines and slopes of the plots. Separate regression analysis was conducted for the data obtained from inside the iron-enriched waters due to a kink after the onset of algal physiological stress confirmed by changes in photochemical quantum efficiency (Fv/Fm).

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During SERIES, the concentrations of silicic acid were almost depleted to zero, yet nitrate was present at 3–4 μmol L–1 by Day 18 of the iron-induced phytoplankton bloom (Boyd et al. 2004). On Day 12, measurements of photosynthetic competence (Fv/Fm) indicated that diatoms were no longer iron replete but were exhibiting algal iron stress. During Days 12–16, the absolute nitrate uptake rates in the surface mixed layer remained at high levels (Marchetti et al. 2006d) (note that Day 1 in Marchetti et al. (2006d) corresponds to Day 0 in this text and that of Boyd et al. (2004)). There was a concurrent, nearly constant consumption in mixed-layer nitrate, but a large increase in the consumption of silicic acid and, hence, an increase in [Si(OH)4]:[NO3] drawdown ratio during Day 12 and Day 18 were observed (Boyd et al. 2005; Marchetti et al. 2006d; Saito et al. 2006b). The drivers of this observed change in [Si(OH)4]:[NO3] stoichiometry could be different from those observed in SEEDS, that is, silicic acid uptake and not nitrate uptake determines the [Si(OH)4]:[NO3] uptake stoichiometry. By using a marine ecosystem model (NEMURO), Takeda et al. (2006) demonstrated that the diatom growth rate is an important factor determining the relative consumption of silicic acid and nitrate following the onset of iron stress, although the potential influence of floristic shifts in the phytoplankton population and decrease in the absolute nitrate uptake rate after Day 18 cannot be ruled out (Marchetti et al. 2006d). With regard to some oceanic pennate diatoms that were dominant during SERIES, Marchetti and Harrison (2007) also found that iron-deficient cells acclimate to low iron concentrations by changing their cell morphology, and that increase in the Si-containing valve surface area relative to the volume of the internal components may influence the cellular Si:N ratio.

During SEEDS, nitrate uptake by the phytoplankton was enhanced by 20 folds after the addition of iron, and macronutrients were initially utilized at a ratio of 15.5 N:1 P:24.6 Si (Kudo et al. 2005). After Day 10, an increase in [Si(OH)4]:[NO3] drawdown ratio was observed, and it was driven by a decrease in nitrate consumption (Fig. 30). At that time, the measured decrease in photosynthetic competence (Fv/Fm) indicated the occurrence of physiological stress on phytoplankton by both low ambient dissolved iron concentration and low light levels in the surface mixed layer due to self-shading by the dense diatom bloom (Tsuda et al. 2003). Differences in the primary production rate and nitrate uptake rate between the top and bottom of the surface mixed layer support the physiological stress by low light availability near the bottom of the euphotic zone (Kudo et al. 2005). On the other hand, gross and net growth rates of phytoplankton remained high in the surface water even after Day 10 (Saito et al. 2005), probably because of the enhanced supply of regenerated iron through active grazing by heterotrophic dinoflagellates (Takeda and Tsuda 2005). Thus, algal iron stress does not appear to be adequately strong to reduce the phytoplankton growth during SEEDS. Although microzooplankton grazing could also regenerate ammonium and have an influence on the nitrate uptake by phytoplankton, both ammonium concentration and its uptake rate by the phytoplankton remained low during SEEDS (Kudo et al. 2005). Most of the grazed phytoplankton were accumulated as a biomass of heterotrophic dinoflagellates (Saito et al. 2005), and ammonium regeneration appeared to be minimum.

An important factor that should be considered in SEEDS is the significant decrease in light availability owing to self-shading by dense phytoplankton cells. Based on the laboratory culture studies, it is known that decrease in light availability can increase the diatom Si:N ratio and [Si(OH)4]:[NO3] uptake ratio (Davis 1976; Saito and Tsuda 2003). However, situation in the natural bloom in the open ocean, where nighttime vertical convection mixes phytoplankton populations acclimated to different light levels, is different from that in the laboratory experiment using diatom cultures acclimated to a different but constant light level.

After the development of dense diatom bloom, cells in the surface layer can actively take up both silicic acid and nitrate, while cells near the bottom of the surface mixed layer may take up nitrate at a lower rate relative to the surface; however, the difference in the uptake of silicic acid between the surface and near the bottom of the surface mixed layer may not be large. This is because the steps of uptake, reduction, and assimilation of nitrate require energy derived from photosynthesis and, therefore, potentially influenced by low light (Vincent 1992), while there is no direct light requirement for the uptake of silicic acid by diatoms, although the cells must be actively growing to take up silicic acid (Blank and Sullivan 1979). Thus, uptake of silicic acid often extends to 1.5–2 times of the maximum depths to which photosynthesis is measurable and also continues through the night in many marine surface waters (e.g., Brzezinski and Nelson 1989; Nelson et al. 1991). Accordingly, during nighttime, vertical convection mixes up together the phytoplankton cells grown under different physiological conditions, and differential uptake between silicic acid and nitrate in the surface mixed layer continues. The vertical convection during nighttime also homogenizes the vertical gradient in nitrate concentration as well as phytoplankton biomass within the surface mixed layer.

By assuming that the uptake of silicic acid by the diatoms takes place within the whole surface mixed layer at a rate that is undiminished from that in the surface, and that the rate of nitrate uptake decreases with depth in proportion to the photosynthetic rate, the [Si(OH)4]:[NO3] uptake ratio integrated within the surface mixed layer was estimated to be about 1.5 during Days 0–7 and about 2.6 during Days 9–13 (Fig. 31). Therefore, most of the observed increase in [Si(OH)4]:[NO3] drawdown ratio in SEEDS bloom could be attributed to the decrease in the nitrate uptake in the surface mixed layer by low light availability due to self-shading effect during the dense diatom biomass period. This mechanism may also partially explain the observed change in [Si(OH)4]:[NO3] drawdown ratio during the coastal diatom blooms in the western subarctic Pacific (Saito et al. 2002). To confirm this "low-light and mixing" hypothesis, diurnal change in nutrient uptake should be determined near the bottom of the surface mixed layer during diatom blooms.


Fig. 31. Estimated uptake rates of silicic acid and nitrate integrated within the surface mixed layer (a) and the ratio of these estimated uptake rates (b) during SEEDS. The following assumptions were made for the estimation of these nutrients' uptake rate: (i) nitrate uptake rate is in proportion to the photosynthetic rate at a 106C:16N ratio; (ii) [Si(OH)4]:[NO3] uptake ratio in the surface water is 1.5; and (iii) silicic acid uptake by diatoms takes place within the whole surface mixed layer at a rate that is undiminished from that in the surface. The photosynthetic rates measured at 6 light depths (100, 35, 25, 10, 6, and 1% of the surface irradiance) during SEEDS were obtained from Kudo et al. (2005). The [Si(OH)4]:[NO3] drawdown ratio (1.5) observed during the early phase of the SEEDS bloom was used for estimating the uptake rate of silicic acid from that of nitrate in the surface water.

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Differences in the response of nutrient consumption upon algal physiological stress between SEEDS and SERIES suggest that different factors influence the bloom toward the end of each bloom. In SEEDS, decrease in light availability and increased grazing by heterotrophic dinoflagellates on the diatoms were important in terminating the bloom, while in SERIES, it was a case of bottom-up control due to algal iron stress, followed by co-limitation by iron and silicic acid. Contrary to the SEEDS and SERIES, the drawdown ratio of [Si(OH)4]:[NO3] decreased at the later phase of the SEEDS-II bloom (Saito et al. 2009). This decrease was due to the increase in the ammonium uptake relative to nitrate in response to the ammonium accumulation by active mesozooplankton grazing during this period (Kudo et al. 2009).

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6. Summary and a view to the future

The idea that the supply of trace element iron may limit the biological productivity of some parts of the ocean has generated a great deal of research and speculation. Over the past several decades, the advances in the knowledge about the physiological and ecological importance of iron for phytoplankton growth as well as the complexity of both the physicochemical iron dynamics in seawater and biological iron uptake strategy have inevitably given rise to a whole new set of questions.

Recent findings on the lateral iron transport from the iron-rich marginal regions to the pelagic waters and its influences on the variation in the surface dissolved iron concentrations highlight the importance of sediment and particulate iron in the subarctic North Pacific. In fact, iron released from the sediments has been considered to have an important contribution to the natural supply of dissolved iron to the surface water at massive bloom sites in the Southern Ocean (Planquette et al. 2007; Blain et al. 2008). There is a relatively consistent and/or seasonal supply of iron through lateral transport from a marginal sea and continental margin as well as vertical mixing during winter in the western (Oyashio) region, while restrictive supply of iron by mesoscale eddy, tidal currents, Ekman transport, and southeastward advection via the Alaska Gyre may produce relatively severe iron limitation in the eastern region. Characteristics and behavior of particulate iron can vary significantly from small colloidal inorganic mineral particles to large organic aggregates, which may include both living and dead components (Hurst and Bruland 2007). In addition, chemical leaching technique to estimate the biologically available (labile) iron is quite different between marine suspended particles and atmospheric dust (Berger et al. 2008; Boyd et al. 2009). Therefore, bioavailability and reactivity of particulate iron cannot be generalized, and elucidation of key processes that regulate the production of suspended iron particles over the continental margin as well as the transformation during lateral transport is a big challenge for both biogeochemists and modelers (Lam et al. 2006; Moore and Braucher 2008; Fiechter et al. 2009). In addition, organic ligands should have interaction with both dissolved and particulate iron, but the measure sources of iron-complexing ligands in the surface/subsurface waters are still unknown. Presence of relatively high concentrations of excess free organic ligands may indicate that formation/release of iron-complexing organic ligands, such as porphyrins, is more active in the western region, where biological processes and cycling are working substantially than the eastern regions. On the other hand, severe iron-limited condition in the eastern region may stimulate bacterial production of siderophore; thus, having the high-affinity iron transport system to take up iron from siderophore-iron complex, which is known to operate in some oceanic centric diatoms, could be one of the strategies for survival, but their growth might be controlled by the absolute amounts of iron in the surface mixed layer. Under such conditions, it is necessary for the phytoplankton to reduce either their cell size (Sunda and Huntsman 1995) or their metabolic requirement for iron (Peers and Price 2006), or both.

In spite of the active researches in the subarctic HNLC waters, significant impacts of Asian dust on the phytoplankton productivity have not been detected, suggesting spatial and temporal mismatch of the dust inputs and biological activities. On the other hand, basin-scale dispersions of volcanic ash stimulated the primary production in nearly the whole area of the subarctic waters in summer (Hamme et al. 2010), but there is limited information on the solubility, bioavailability, and toxicity of the ash particles. Physicochemical processing of aerosol particles within acidic fogs may also play an important role over the foggy subarctic waters during summer. These new findings give us a useful hint to make new advances in understanding the iron biogeochemical cycles in the subarctic North Pacific.

Laboratory and onboard culture experiments of subarctic phytoplankton indicate severe limitation of their growth at subnanomolar iron concentration levels usually found in the HNLC oceanic waters. Satisfaction of algal demands for both light and iron is a key for phytoplankton blooming in the HNLC waters, and it seems to be interesting to compare the range of fluctuation in iron requirement under saturating and limiting light conditions between pennate and centric diatoms for a better understanding of the influence of iron and light co-limitation on the growth competition between them within distinct phytoplankton communities in the eastern and western subarctic gyres. To address some of these important questions, it appears that the experimental tools associated with recent advances in genomic science, such as genomic microarray experiment (Allen et al. 2008), should be practically introduced into the field of oceanography in addition to the traditional laboratory/onboard culture approaches and advanced trace-metal clean observations. Surprisingly, the community half-saturation constant for growth with respect to iron was similar between the western and eastern gyres. However, differences in the iron supply process and its availability in these two gyres seem to have developed unique populations dominated by the centric diatoms in the west and the pennate diatoms in the east. It is essential to evaluate iron transport processes to the open ocean by considering the time-scales needed for phytoplankton blooms. It can be noted that the lateral transport of particulate matters from continental margins may also supply seeds or cyst of coastal phytoplankton species to the pelagic waters.

Three open-ocean iron-enrichment experiments conducted in the Western Subarctic Gyre and the Alaska Gyre have renewed the appreciation about the importance of physical processes, such as fluctuation of vertical mixing layer, horizontal advection/diffusion, and presence of mesoscale eddies as well as the strong coupling within pelagic food web through both top-down and bottom-up controls. For example, the observed increases of large, long setae, or heavily silicified diatoms, such as Chaetoceros convolutes, during the high grazing pressure period in the SEEDS-II experiment have been attributed as a result of anti-predation from copepods (Tsuda et al. 2009), and it strongly suggests that both moderate iron supply and active mesozooplankton grazing may lead the phytoplankton assemblage to the dominance of large centric diatoms in the western region. On the other hand, low Fe:C ratio and high iron usage efficiency would be advantageous for coccolithophorids and pennate diatoms to survive in consistently low iron conditions in the eastern region with the additional prominence for pennates that have an extensive iron-storage ability (Marchetti et al. 2009) for sporadic iron supply from the atmosphere or the continental shelf sediments. The difficulties in studying the processes spread over both the surface and subsurface systems, such as the biological carbon pump, on a spatiotemporal scale large enough to cover a bloom event by usual shipboard observations have also been realized. One of the interesting, yet speculative, subjects is the sources of iron-complexing organic ligands and the biological availability of organically complexed iron in the water column. It appears to be difficult to reproduce the dynamic processes related to these subjects in a small carboy on a research vessel. The Lagrangian coordinate system established for mesoscale enrichment experiments could be applicable for investigating such dynamic processes by tracking a mesoscale eddy or spring bloom water mass.

Besides the continuous efforts to understand the differences between the western and the eastern gyres, new studies focusing on the central regions of the subarctic North Pacific would be attractive, especially at the waters affected by Alaskan Stream (Ueno et al. 2009), where iron, macronutrients, and plankton could be transported in both east-west and north-south directions along the Aleutian Islands, and at the Subarctic Boundary, where the HNLC waters encounter the subtropical surface waters containing iron to some degree. These studies are certainly needed to draw a whole picture of the iron-phytoplankton interactions in the subarctic North Pacific.

Acknowledgments

The author is thankful to J. Nishioka and Y. Kondo for their supports on the analyses of iron and iron-complexing organic ligands. The author is also grateful to all the scientists and ship crew members of the OPES, SEEDS, SERIES, and SEEDS-II projects. This work was supported in part by grants from the Central Research Institute of Electric Power Industry, the Ministry of Education, Science and Culture, the Global Environmental Research Fund from the Ministry of Environment, and the Fisheries Agency. The comments and suggestions of the anonymous reviewers are also gratefully acknowledged.

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List of Figures

Fig. 1. General circulation in the Subarctic North Pacific showing the Alaska Gyre and Western Subarctic Gyre. Double arrows are intense boundary currents. The Subarctic Boundary separates the subarctic Pacific region to the north from the subtropical Pacific region to the south. Reprinted from Prog. Oceanogr., 43, Harrison et al., Comparison of factors controlling phytoplankton productivity in the NE and NW subarctic Pacific Gyres, 205–234, © 1999, with permission from Elsevier.

Fig. 2. Vertical distribution of size-fractionated iron in the western subarctic North Pacific (WSNP) and the eastern subarctic North Pacific (ESNP). (a) The data were collected at Oyashio region (49°N; 157°30'E) in May 2000; observed surface mixed layer was 30 m. (b) The data were collected at St. KNOT (44°N; 155°E) in May 2000; observed surface mixed layer was 10 m. (c) The data were collected at Ocean Station P in September 1998; observed surface mixed layer was 40 m. Nishioka et al., Size-fractionated iron distributions and iron-limitation processes in the subarctic NW Pacific, Geophys. Res. Lett., 30(14), 1730, © 2003 American Geophysical Union. Reproduced by permission of American Geophysical Union.

Fig. 3. (a) Vertical profile of dissolved iron (D-Fe), iron-complexing organic ligand concentrations, and (b) conditional stability constant with respect to inorganic Fe' (K'Fe'L) measured at 50°N, 160°W (Station 13, KH-05-2 cruise) in the eastern subarctic Pacific. The error bars show the analytical errors (Kondo and Takeda, unpublished).

Fig. 4. Comparison of iron-complexing organic ligands with dissolved iron concentrations in (a) surface water (10 m) and (b) deep water (75–5000 m) measured in the subarctic North Pacific. The dotted line shows 1:1 ratio (Kondo and Takeda, unpublished).

Fig. 5. Vertical distributions of the dissolved iron concentrations at Ocean Station P in September 1997 and February 1999. The data reported by Martin et al. (1989) at Station T-7 were also plotted for comparison. Reprinted from Mar. Chem., 74, Nishioka et al., Size-fractionated iron concentrations in the northeast Pacific Ocean: Distribution of soluble and small colloidal iron, 157–179, © 2001, with permission from Elsevier.

Fig. 6. Temporal change in iron distributions of each size fraction at Ocean Station P in September 1997, June 1998, September 1998, and February 1999. (a) Soluble Fe, (b) colloidal Fe, and (c) labile particulate Fe. Reprinted from Mar. Chem., 74, Nishioka et al., Size-fractionated iron concentrations in the northeast Pacific Ocean: Distribution of soluble and small colloidal iron, 157–179, © 2001, with permission from Elsevier.

Fig. 7. Comparison of colloidal (200 kDa-0.2 μm) iron with phosphate concentrations at Ocean Station P in September 1997 and February 1999. Reprinted from Mar. Chem., 74, Nishioka et al., Size-fractionated iron concentrations in the northeast Pacific Ocean: Distribution of soluble and small colloidal iron, 157–179, © 2001, with permission from Elsevier.

Fig. 8. Sampling stations for observational studies on iron distributions and for both onboard and in situ iron-enrichment experiments. The arrows indicate a schematic image of the surface water currents.

Fig. 9. (a) Vertical profiles of dissolved iron (D-Fe), total iron (T-Fe), salinity, and nitrate + nitrite concentrations from January to May at Station A7 (41°30'N, 145°30'E; upstream of Oyashio region) and (b) from March to May at Station B9 (44°N, 155°E; downstream of Oyashio region) in 2003. Nishioka et al., Iron supply to the western subarctic Pacific: importance of iron export from the Sea of Okhotsk, J. Geophys. Res., 112, C10012, © 2007 American Geophysical Union. Reproduced by permission of American Geophysical Union.

Fig. 10. Seasonal variations in (a) sea-surface dissolved iron concentrations (average in surface-mixed layer); (b) nitrate + nitrite concentrations (average in surface-mixed layer); (c) surface-mixed layer depths (MLD); and (d) chlorophyll a concentrations (average in surface-mixed layer) from January to the end of May 2003 along the "A-line". Nishioka et al., Iron supply to the western subarctic Pacific: importance of iron export from the Sea of Okhotsk, J. Geophys. Res., 112, C10012, © 2007 American Geophysical Union. Reproduced by permission of American Geophysical Union.

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Fig. 11. Schematic of the iron supply process in the subarctic NW Pacific. Water ventilation processes in this region control the transport of dissolved and particulate iron through the intermediate water layer from the continental shelf of the Sea of Okhotsk to the wide area of the western subarctic Pacific. Nishioka et al., Iron supply to the western subarctic Pacific: importance of iron export from the Sea of Okhotsk, J. Geophys. Res., 112, C10012, © 2007 American Geophysical Union. Reproduced by permission of American Geophysical Union.

Fig. 12. A longitudinal section of dissolved iron and total iron profile in the North Pacific along 165°E. A density range of 26.6–27.5σθ is located on each figure. Nishioka et al., Iron supply to the western subarctic Pacific: importance of iron export from the Sea of Okhotsk, J. Geophys. Res., 112, C10012, © 2007 American Geophysical Union. Reproduced by permission of American Geophysical Union.

Fig. 13. Relationships between total dissolved iron concentration and specific growth rate (a, b) and final cell yield measured at the stationary phase (c, d) for Emiliania huxleyi (○), Gephyrocapsa oceanica (■), Thalassiosira sp. (▲), Actinocyclus sp. (□), and Chaetoceros sp. (●) isolated from NW Pacific.

Fig. 14. Changes in the concentrations of size-fractionated chlorophyll a (a, 0.7–5 μm; b, >5 μm), silicic acid (c), phosphate (d), and nitrate (e) in control and experimental bottles enriched with various amounts of iron. The experiments were conducted at Ocean Station P (50°N; 145°W) in September 1998.

Fig. 15. Specific growth rates of phytoplankton assemblage, calculated using concentrations of chlorophyll a in two size classes (0.7–5 μm [] and >5 μm []), plotted against initial dissolved iron concentrations. The data were from the bottle incubation experiments at Ocean Station P (50°N; 145°W) in September 1998. The open triangles (0.7–5 μm [] and >5 μm []) represent data from the incubation bottles treated with 100 nM desferrioxamine B.

Fig. 16. (a) Changes in concentrations of size-fractionated chlorophyll a in control and experimental bottles enriched with various amounts of iron, measured at the beginning of the experiment (initial) and after 4 and 6 days of incubation. (b) Relationships between initial dissolved iron concentration and specific growth rates of phytoplankton assemblage, calculated using size-fractionated chlorophyll a data. The onboard incubation was conducted during the SEEDS-II experiment using iron-fertilized surface water on Day 17.

Fig. 17. Time-course change in chlorophyll a concentrations in three size fractions in controls and Fe-enriched bottles during on-deck incubation at Station A in the subarctic NW Pacific. Dissolved iron concentration in the control bottles was 0.22 nM. Values are the mean of duplicate samples.

Fig. 18. Changes in dissolved iron concentration in incubation bottles enriched with iron (●) and in controls (○) during on-deck incubation at Station A in the subarctic NW Pacific. The error bars are the standard deviation of the duplicate samples.

Fig. 19. Time course of nitrate (a), phosphate (b), and silicic acid (c) concentrations in controls (●) and incubation bottles enriched with iron (□) or DFB (△) during on-deck incubation at Station A4 in May 2003. The values are the means of triplicate samples, and the error bars are the standard deviation. Reprinted from PhD Thesis, Yoshiko Kondo, Dynamics of the organic Fe complexing ligands and phytoplankton in the Pacific Ocean, 256 pp., © 2007.

Fig. 20. Changes in chlorophyll a concentration in two size fractions in controls (a) and incubation bottles enriched with iron (b) or DFB (c) during on-deck incubation at Station A4 in May 2003. The values are the means of triplicate samples, and the error bars are the standard deviation. Reprinted from PhD Thesis, Yoshiko Kondo, Dynamics of the organic Fe complexing ligands and phytoplankton in the Pacific Ocean, 256 pp., © 2007.

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Fig. 21. Changes in chlorophyll a concentration in two size fractions in controls (a, d) and incubation bottles enriched with iron (b, e) or DFB (c, f) during on-deck incubation at stations A7 (a–c) and A17 (d–f) in May 2003. The values are the means of triplicate samples. Reprinted from PhD Thesis, Yoshiko Kondo, Dynamics of the organic Fe complexing ligands and phytoplankton in the Pacific Ocean, 256 pp., © 2007.

Fig. 22. Time course of nitrate (a, d), phosphate (b, e), and silicic acid (c, f) concentrations in controls (●) and incubation bottles enriched with iron (□) or DFB (▲) during on-deck incubation at Station A7 (a–c) and A17 (d–f) in May 2003. The values are the means of triplicate samples. Reprinted from PhD Thesis, Yoshiko Kondo, Dynamics of the organic Fe complexing ligands and phytoplankton in the Pacific Ocean, 256 pp., © 2007.

Fig. 23. Changes in chlorophyll a concentration in two size fractions in controls (low temperature, low light and low-iron) and incubation bottles treated with high temperature, high light, high iron, and combinations of those (T+L, high temperature and high light; T+L+Fe, high temperature and high light with iron addition). The experiments were conducted onboard using surface water collected at Ocean Station P in February 1998. See Table 6 for the detailed experimental conditions.

Fig. 24. Location of the experimental sites for SEEDS, SEEDS-II in the Western Subarctic Gyre, and SERIES near the Ocean Station P (P26) in the Alaska Gyre.

Fig. 25. A comparison of changes in surface dissolved iron (a) and chlorophyll a (b) concentrations in the surface water during the time course of SEEDS, SERIES, and SEEDS-II. The data are from Tsuda et al. (2003), Boyd et al. (2004), Wong et al. (2006b), Tsuda et al. (2007), and Nishioka et al. (2009).

Fig. 26. A comparison of changes in surface concentrations of nitrate (a) and silicic acid (b) during the time course of SEEDS, SERIES, and SEEDS-II. The data are from Tsuda et al. (2003), Boyd et al. (2004), and Tsuda et al. (2007).

Fig. 27. A comparison of changes in contribution of diatoms to total chlorophyll a biomass during the time course of SEEDS, SERIES, and SEEDS-II. The data are from Suzuki et al. (2005), Wong and Crawford (2006), and Suzuki et al. (2009).

Fig. 28. A comparison of time-series variations of POC-sinking flux collected at 40- or 50-m sediment traps inside and outside of the iron-fertilized patch during SEEDS, SERIES, and SEEDS-II. The data are from Tsuda et al. (2003), Boyd et al. (2004), and Aramaki et al. (2009).

Fig. 29. Time-series measurements of the concentrations of iron-complexing organic ligands and dissolved iron in the surface water during SEEDS (a, Takeda and Kondo, unpublished) and SEEDS-II (b, Kondo et al. 2008). The error bars show the analytical errors.

Fig. 30. Relationship between nitrate and silicic acid concentrations at 5- to 10-m depth inside and outside of the iron-enriched waters measured during SEEDS and SERIES. The lines and Si/N ratio designate the linear regression lines and slopes of the plots. Separate regression analysis was conducted for the data obtained from inside the iron-enriched waters due to a kink after the onset of algal physiological stress confirmed by changes in photochemical quantum efficiency (Fv/Fm).

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Fig. 31. Estimated uptake rates of silicic acid and nitrate integrated within the surface mixed layer (a) and the ratio of these estimated uptake rates (b) during SEEDS. The following assumptions were made for the estimation of these nutrients' uptake rate: (i) nitrate uptake rate is in proportion to the photosynthetic rate at a 106C:16N ratio; (ii) [Si(OH)4]:[NO3] uptake ratio in the surface water is 1.5; and (iii) silicic acid uptake by diatoms takes place within the whole surface mixed layer at a rate that is undiminished from that in the surface. The photosynthetic rates measured at 6 light depths (100, 35, 25, 10, 6, and 1% of the surface irradiance) during SEEDS were obtained from Kudo et al. (2005). The [Si(OH)4]:[NO3] drawdown ratio (1.5) observed during the early phase of the SEEDS bloom was used for estimating the uptake rate of silicic acid from that of nitrate in the surface water.

Table 1. Comparison of iron concentration in labile particulate and dissolved fractions in the surface mixed layer between the Western and Eastern Subarctic North Pacific. Nishioka et al., Size-fractionated iron distributions and iron-limitation processes in the subarctic NW Pacific, Geophys. Res. Lett., 30(14), 1730, © 2003 American Geophysical Union. Reproduced by permission of American Geophysical Union. aThe data are shown as the average value ± the standard deviation calculated for labile particulate and dissolved iron concentrations in the surface mixed layer at stations in Oyashio region (May–June 2000). bIn WSNP region (May–June 2000, and July–Aug. 2001). cIn ESNP (Sep. 1997 and Feb., June, and Sep. 1998).

Table 2. Phytoplankton strains isolated from the NW Pacific for laboratory culture experiments.

Table 3. Cellular iron to carbon ratio (Fe/C) and iron use efficiency of the NE Pacific phytoplankton under iron stress. Estimation was based on the measured POC value at the stationary phase of the growth and by assuming that all the iron was taken up under the iron-limited condition.

Table 4. Comparison of the half saturation constant for iron (Ks) for the NW Pacific phytoplankton with the reported value for other oceanic strains. *The concentration of dissolved inorganic iron (Fe') was used instead of the total dissolved iron for the estimation of Ks. (a) Timmermans et al. (2001); (b) Sunda and Huntsman (1995); (c) Kudo and Harrison (1997).

Table 5. Comparison of the half saturation constant for iron (Ks) for subarctic phytoplankton assemblage in two size classes. *The total (unfiltered) acid-labile iron concentration in the control bottles was measured on the last day of incubation. (a) Kudo et al. (2006); (b) Noiri et al. (2005). Iron was added as a ferric chloride solution in these experiments as well as in this study.

Table 6. Experimental conditions for light, temperature, and iron treatments and control during the bottle incubation experiments at Ocean Station P (50°N, 145°W) in February 1998. T+L — high-temperature and high-light; T+L+Fe — high-temperature and high-light with iron addition.

Table 7. Net growth rates of phytoplankton biomass during the iron enrichment experiments in the subarctic NW Pacific (Station A). *The rate difference between the control and iron treatment is statistically significant (P < 0.01).

Table 8. Dissolved iron concentration at the initial time point (Day 0) and final time point (Day 5) in the controls and iron-enriched samples during the iron enrichment experiments in the subarctic NW Pacific (Station A). Decrease percentage is calculated as 100 × (initial Fe – final Fe)/initial Fe. The ratios of Fe decrease to particulate organic carbon increase (Fe:POC) in the control and iron-enriched samples are also shown. The values are the mean of duplicates and the standard deviation.

Table 9. Initial (Day 0) and final (Day 5) concentrations of POC and Chl a and the ratio of POC/Chl a for the controls and iron-enriched samples during the iron enrichment experiments in the subarctic NW Pacific (Station A). The values are the means of duplicate samples and the standard deviation.

Table 10. Initial conditions at Stations A4, A7, and A17 in the Oyashio region in May 2003. Reprinted from PhD Thesis, Yoshiko Kondo, Dynamics of the organic Fe complexing ligands and phytoplankton in the Pacific Ocean, 256 pp., © 2007.

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Table 11. Initial and final cell densities (cells ml–1) for each size class of phytoplankton and microzooplankton during the bottle incubation experiments at Ocean Station P (50°N, 145°W) in February 1998.

Table 12. Comparison of the initial conditions in the surface water among SEEDS, SERIES, and SEEDS-II. *Fv/Fm: photochemical quantum efficiencies of algal photosystem II. Data are from Takeda and Tsuda (2005), Boyd et al. (2005), and Tsuda et al. (2007).

Table 13. Times and quantities comprising the iron infusion in SEEDS, SERIES, and SEEDS-II experiments.

Table 14. A comparison of the fugacity of carbon dioxide (fCO2) in the unfertilized ambient water and the minimum value observed during the SEEDS, SERIES, and SEEDS-II experiments. Data are obtained from Tsuda et al. (2003, SEEDS), Wong et al. (2006a, SERIES), and Tsuda et al. (2007, SEEDS-II).

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