Biogeochemistry of a Changing Arctic Ocean

Kevin R. Arrigo

Department of Environmental Earth System Science, Stanford University

1. Introduction

The Arctic Ocean is currently in the forefront of climate change caused by both natural and anthropogenic factors. Since 1950, the mean annual air temperature has increased by 2-3°C and by 4°C in winter (Chapman and Walsh 2003), resulting in markedly longer summers (Smith 1998). Temperature is projected to increase by an additional 4-5°C by the end of the 21st century (ACIA 2005). In conjunction with these higher temperatures, sea ice cover in the Arctic Ocean has been contracting over the past three decades, with dramatic reductions in recent years (Levi 2000, Parkinson 2000). Changes in sea ice cover also include an increase in the length of the ice melt season (Smith 1998, Rigor et al. 2002, Serreze et al. 2007, Comiso et al. 2008) and a decrease in ice thickness over the central Arctic Ocean (Rothrock et al. 1999). The result is greater open water area and enhanced shelf break upwelling, the latter of which is expected to increase the input of nutrients from offshore waters to shallower shelves (ACIA 2005). While a reduction in sea ice should favor the growth of phytoplankton and increase the net air-to-sea flux of CO2 (Anderson and Kaltin 2001, Bates et al. 2006), it also will reduce the amount of production contributed by algae growing within the sea ice (Subba Rao and Platt 1984, Legendre et al. 1992, Gosselin et al. 1997), although sea ice communities generally account for a relatively small fraction of total primary production in Arctic waters.

Recent declines in Arctic sea ice cover have been attributed to a combination of factors, including increased advection of warm water into the Arctic Ocean, both from the Atlantic through the eastern Fram Strait and the Barents Sea (Steele and Boyd 1998, Dickson et al. 2000) and from the Pacific in the form of relatively warm Pacific Surface Water (Maslowski et al. 2001, Shimada et al. 2006), atmospheric circulation patterns that favor advection of sea ice out of the Arctic Ocean through Fram Strait (Rigor and Wallace 2004, Maslanik et al. 2007, Serreze et al. 2007), and increased Arctic temperatures (Rothrock and Zhang 2005, Lindsay and Zhang 2005). As thick multi-year sea ice has been replaced by a thinner sea ice cover, the ice pack is more easily melted, either by surface heating or advection of warm waters into the Arctic Ocean. Reduced sea-ice extent decreases surface albedo, allowing more shortwave radiation to penetrate the ocean surface, contributing to additional ocean heat content and thus creating a positive feedback mechanism that inhibits ice growth in winter and accelerates its loss in spring and summer (Perovich et al. 2007).

Reported impacts of reduced Arctic sea ice extent include increasing autumn and winter temperatures, stronger wave activity and intensified coastal erosion (Serreze et al. 2007), disruption of thermohaline circulation (Peterson et al. 2006), impairment of traditional hunting practices (Huntington and Fox 2005), and improved navigation through the newly opened Northwest Passage. Thus far, the impact that reduced sea ice cover will have on pan-Arctic marine primary production has received little attention (Pabi et al. 2008). Although Arctic sea ice itself can be biologically productive (Gosselin et al. 1997, Mock and Gradinger 1999), occasionally supporting large populations of diatoms and other primary producers, areal rates of CO2-fixation in sea ice habitats tend to be much lower than rates found in the adjacent ice-free ocean. Therefore, a loss of Arctic sea ice might be expected to increase the area favorable for phytoplankton growth and enhance the productivity of the Arctic Ocean.
Unlike other oceans, the Arctic Ocean is almost completely landlocked, except for the very shallow Bering Strait (~50 m) that connects it to the Pacific Ocean, and the Fram Strait and Canadian Archipelago that allow exchange with the Atlantic Ocean. Associated with the extensive land margin is a broad continental shelf, 5  106 km2 in area, and comprising about 53% of the total Arctic Ocean. This is much higher than the 9.1 -17.7% characteristic of continental shelves in other oceans of the world (Menard and Smith 1966, Jakobsson et al. 2003). Ice-free continental shelves, such as those found in parts of the Chukchi Sea, often experience intense seasonal blooms of phytoplankton owing to their favorable nutrient and light conditions (Hill and Cota 2005).

Another unique feature of the Arctic Ocean is the large amount of riverine discharge it receives (~4000 km3 yr-1) (Shiklomanov 2000, Carmack and Macdonald 2002), arising from both large rivers, like the Ob, Lena, Yenisey, and Mackenzie, and numerous smaller ones in both the Amerasian and Eurasian sectors. This large freshwater input affects both the salinity and nutrient concentration of the Arctic Ocean. Furthermore, it is predicted that precipitation in a warming climate will increase significantly (IPCC 2006), thereby enhancing the already enormous fluxes of riverine sediment discharge (670 Mt yr-1) and organic carbon (12.6 Mt yr-1) from the land to the Arctic Ocean (Macdonald et al. 1998), both of which will impact nutrient and light availability and hence, phytoplankton growth.

The circulation of the Arctic Ocean is comprised of both low salinity (<33) and nutrient-rich Pacific Ocean water and relatively nutrient-poor and more saline (~34.8) Atlantic Ocean water (Maslowski et al. 2004). The denser Atlantic water is distributed via counterclockwise currents along the continental slope at the basin margins. The relatively less saline and warmer Pacific water enters the Arctic Basin through the Bering Strait between Cape Dezhnev and west Alaska, and exits through the Canadian Archipelago, the Fram Strait and the Nares Strait. Historically, the front separating the Atlantic and Pacific water has been located over the Lomonosov Ridge, but recently this front appears to have moved closer to the Alpha-Mendeleev Ridge. This shift in the location of the front has lead to the displacement of a large quantity of Pacific water that has been replaced by nutrient-poor water from the Atlantic (Macdonald 1996), potentially reducing the amount of nutrients available for phytoplankton growth. The nutrient-rich water from the Pacific Ocean is generally restricted to the Chukchi Sea and the Amerasian Basin (Carmack et al. 1997).

Surface concentrations of nitrate, phosphate and silicic acid in Arctic waters approach detection limits after the spring bloom (Sakshaug 2003), suggesting that annual primary production is generally controlled by nutrient availability. The nitrate to phosphate ratio in these waters ranges from 11 to 16 (mol:mol) (Sakshaug 2003), suggesting that much of the Arctic Ocean is nitrogen-limited (assuming that phytoplankton require nitrogen and phosphorus at the Redfield ratio of 16:1). Phosphorus limitation of phytoplankton is more likely in waters with a salinity of <25 (Sakshaug et al. 1983) due to low phosphate content of river waters that are otherwise rich in nitrate. The silicic acid to nitrate molar ratio is spatially variable, ranging from a high of 1.9-2.4 in the Chukchi Sea and Eastern Canadian Arctic to a low of 0.31 in the Eurasian basin (Codispoti 1979, Harrison and Cota 1991, Sakshaug 2003).

Finally, sea ice dynamics are integral to the regulation of primary production in much of the Arctic Ocean. In winter, brine rejection due to ice formation destabilizes the mixed layer, leading to deep vertical mixing and replenishment of surface nutrient inventories. In spring, melting of ice results in strong surface ocean stratification, exposing the nutrient-rich waters to a light regime suitable for phytoplankton growth. The resulting spring ice edge bloom forms a significant component of the annual primary production (Niebauer et al. 1990, Falk-Petersen et al. 2000).
The Arctic Ocean is an ideal location in which to study the ongoing alterations in a marine ecosystem that is already being heavily impacted by changes in climate. These changes are likely to modify the physics, biogeochemistry, and ecology of this unique environment in ways that are not yet understood. Satellite remote sensing has provided some insight into these changes (Pabi et al. 2008, Arrigo et al. 2008], but these satellite-based observations have been limited by a lack of both ground-truth data and ancillary information that could broaden the scope of information that satellite sensors currently provide. For example, we know that the large rivers flowing into the Arctic Ocean contribute large amounts of chromophoric dissolved organic matter (CDOM). This CDOM absorbs strongly in the blue region of the spectrum and interferes with our ability to accurately retrieve chlorophyll a concentrations from space. However, CDOM also increases energy absorption within the mixed layer and hastens the melting of sea ice (Hill 2008), and is photodegraded in surface waters and becomes a significant source of CO2. Therefore, understanding the dynamics of CDOM in the Arctic will reap a variety of benefits in topics as diverse as marine bio-optics, carbon cycling, and the Arctic energy balance.

2. Key Scientific Issues

2.1. Land-Ocean Interactions (changes in runoff)

2.1.1. Increases in sediment discharge

In order to predict changes in the biological productivity of the Arctic coastal ecosystem, an accurate understanding of the effects of climate change on river sediment fluxes is needed. It has been estimated that the total suspended matter flux to the Arctic (227 x 106 tons yr-1) is very low, only about 1% of the global flux, due to unusually low concentrations of suspended matter that are an order of magnitude below the global mean. Suspended sediment concentrations in Arctic rivers are low due to thin weathering crusts on the Arctic watersheds, low precipitation, extensive permafrost, low temperatures, large areas of swamps and lakes and a low level of human activity (Syvitski 2002, Gordeeva 2006). Models suggest, however, that sediment flux could increase by 30% for every 2°C of warming, and that a 10% increase in sediment load would be associated with a 20% increase in water discharge. Based on these results, the sediment flux from the six largest Arctic rivers (Yenisey, Lena, Ob, Pechora, Kolyma and Severnaya Dvina ) is predicted to increase by 30-122% by 2100 (Syvitski 2002, Morehead et al. 2003, Gordeeva 2006)

2.1.2. Increases in CDOM

Arctic rivers transport large quantities of dissolved organic carbon (DOC) to the Arctic Ocean. It is widely considered that most of this DOC is refractory and of minor significance to the biogeochemistry of the Arctic Ocean. However, recent evidence suggests that a large fraction of the DOC transported in rivers in the spring is indeed labile (Holmes et al. 2008), suggesting that inputs of DOC to the Arctic Ocean may have a much larger influence on coastal ocean biogeochemistry than previously realized. A substantial fraction of this DOC consists of chromophoric dissolved organic matter (CDOM). This CDOM, along with increased turbidity from river discharge, has been shown to have a marked impact on the attenuation of photosynthetically available radiation (PAR) in the water column, reducing depth-integrated rates of primary production (Retamal et al. 2008). CDOM has an even stronger impact on attenuation of UV, and it has been shown that removal of harmful UV radiation more than offsets the decrease in phytoplankton productivity due to the reduction in PAR (Arrigo and Brown 1996). Furthermore, it has been shown that shifts in CDOM loading (e.g., through climate change, land-use practices, or changes in ocean circulation) can cause variations in biological UV exposure that are of much greater magnitude than those induced by changes in stratospheric ozone (Gibson et al. 2000).

The input of CDOM into the Arctic from its large river system can be quite variable in terms of their concentrations and optical characteristics. For example, the Great Whale River in the eastern Canadian Low Arctic and in the Mackenzie River in the western Canadian High Arctic contained high concentrations of both DOC (3 and 6 mg DOC l-1, respectively) and CDOM (absorption at 320 nm of 11 and 14 m-1, respectively). There were pronounced differences in the CDOM characteristics of these two rivers, with the Mackenzie River being relatively depleted in humic materials. In addition there was a marked decrease in the relative importance of fulvic and humic acid materials between the Mackenzie River and the nearshore coastal environment, in contrast to the much smaller changes observed within the freshwater-saltwater transition zone of the Great Whale River plume (Retamal et al. 2007). Optical differences between marine and terrestrially-derived CDOM can also be large. Analysis of the spectral slope (S) for CDOM absorption in the North Sea and Greenland Sea show marked differences between offshore and coastal waters (Stedmon et al. 2001). These authors suggest that it may be possible to distinguish between terrestrial and marine CDOM in oceanic environments using optical measurements only. Given these observations, models of Arctic continental shelf responses to present and future climate regimes will need to consider these striking regional differences in the organic matter content, biogeochemistry and optics between waters from different catchments and different inshore hydrodynamic regimes (Retamal et al. 2007).

2.1.3. Decreases in salinity

Increased river flow will also impact the freshwater balance of the Arctic, a major controlling factor on thermohaline circulation in the North Atlantic Ocean. In response to the projected 1.4-5.8°C increase in atmospheric temperatures by the year 2100 (IPCC 2007), river discharge into the Arctic Ocean is projected to increase by 0.007 Sv per degree Celsius of temperature increase (Peterson et al. 2008). Discharge from the six largest Eurasian Arctic rivers alone is expected to increase by 18-70% over present conditions. This is comparable to the increase of 35% that has been predicted by the NASA Goddard Institute for Space Studies model in response to a 4°C increase in global temperatures (Van Blarcum et al. 1995).
Changes in river discharge of this magnitude are potentially important with respect to NADW formation. It has been estimated that an increase in the flux of freshwater of 0.06-0.15 Sv would seriously hamper the formation of NADW (Rahmsdorf 2002). Because changes in Arctic river discharge have the potential to impact Atlantic thermohaline circulation over the coming century, it is essential that we now to increase our understanding of the coupled land, ocean, and atmospheric components of the Arctic hydrologic cycle (Peterson et al. 2008).

2.2. Changes in sea ice cover and primary production

Between the late 1970’s and the early part of the 21st century, the extent of Arctic Ocean sea ice cover has declined during all months of the year, with the largest declines reported in the boreal summer months, particularly in September (8.6±2.9% per decade) (Serreze et al. 2007). The loss of Arctic sea ice has accelerated since 2002, with large winter losses reported in 2005 and 2006, a season that usually exhibits little interannual variability (Comiso 2006). Recent results from an ensemble of 11 models used in the International Panel on Climate Change Fourth Assessment Report suggested that there was a high probability of a 40% reduction of summer sea ice extent (relative to the 1979-1999 mean) in the Arctic by the year 2050 (Overland and Wang 2007). However, these predictions were made prior to the summer of 2007, which experienced by far the lowest sea ice cover ever recorded and the largest single year drop in minimum sea ice extent, with a summer minimum that was an unprecedented 23% below the previous low value observed in September 2005 and 39% below the 1979-2000 September mean (NSIDC 2007). Since 1998, open water area in the Arctic has increased at the rate of 0.07  106 km2 yr-1, with the greatest increases in the Barents, Kara and Siberian sectors, particularly over the continental shelf (Pabi et al. in press).

Results from Pabi et al. (in press) show that since 1998, open water area in the Arctic has increased at the rate of 0.07  106 km2 yr-1, with the greatest increases in the Barents, Kara and Siberian sectors, particularly over the continental shelf. They showed that although pan-Arctic primary production averaged 419±33 Tg C yr-1 during 1998-2006, recent increases in open water area have lead to higher rates of annual production, which reached a 10-year peak in 2007 (Arrigo et al. in press). Annual production was roughly equally distributed between pelagic waters (less productive but greater area) and waters located over the continental shelf (more productive but smaller area).

Interannual differences are most tightly linked to changes in sea ice extent, with changes in sea surface temperature (related to the Arctic Oscillation) and incident irradiance playing minor roles. As a result of the progressive loss of sea ice in recent years, annual pan-Arctic primary production increased of by an average of 27.5 Tg C yr-1 each year between 2003 and 2007, with annual production in 2007 (513 Tg C yr-1) exceeding the 1998-2002 mean (416 Tg C yr-1) by 23% (Arrigo et al. in press). Despite its generally low rates of primary productivity, the Siberian sector experienced the largest increase in 2007, with an annual rate (47 Tg C yr-1) that was >3-fold higher than the mean for 1998-2002. Similarly, annual production in both the Laptev and Chukchi sectors was substantially higher (65%) in 2007 than it was in 1998-2002. Changes in annual production in the other Arctic sectors were somewhat smaller, ranging from a 33% increase in the Kara sector in 2007 to a slight decrease (4%) in the Greenland sector (Arrigo et al. in press).

2.2.1. Changes in nutrient inventories

Given that surface nutrients in the Arctic are generally low, it is possible that future increases in production resulting from decreased sea ice extent and a longer phytoplankton growing season will slow as surface nutrient inventories are exhausted. This could reduce primary productivity in waters downstream of the Arctic, such as in the western north Atlantic. On the other hand, nitrate concentrations in subsurface Arctic waters are relatively abundant (approximately 5 µM and 5-15 µM at 50 m and 100 m, respectively). Currently, these nutrients seldom reach the surface due in part to the presence of a cold halocline layer that resides at a depth of 50-100 m and separates deeper high-nutrient waters from the surface layer (Aagaard et al. 1981). However, as multi-year ice continues to retreat from the continental shelf, where most of the primary production currently takes place, wind-driven shelf-break upwelling is likely to be initiated (Carmack and Chapman 2003), increasing the rate of nutrient upwelling onto the shelves and fueling additional increases in productivity.

2.2.2. Changes in oxygen and denitrification

The export of organic matter in the Arctic depends on the structure of the pelagic ecosystem, with a sea ice algae-benthos system exporting carbon more efficiently than an open water phytoplankton-zooplankton dominated system (Piepenberg 2005). However, within the open water ecosystem, higher Arctic primary production and the associated increase in the flux of organic matter to continental shelf sediments also could have important impacts on ocean biogeochemistry. Shelves in the Bering and Chukchi Sea currently are sites of enhanced denitrification (Tanaka et al. 2004, Devol et al. 1997). Further increases in export on the shelves could decrease oxygen concentrations and raise rates of sediment and water column denitrification, resulting in larger losses of fixed inorganic nitrogen and advection of even more excess phosphorus from the Arctic into the Atlantic. Currently, this loss of fixed nitrogen in the Arctic is compensated for by increased fixation of atmospheric nitrogen (N2) in the north Atlantic (Yamamoto-Kawai et al. 2006). However, while N2-fixation is favored in waters with a low nitrogen:phosphorus ratio (Tyrrell 1999), this process also requires high iron concentrations (Zehr et al. 1993). Whether future iron fluxes into the north Atlantic will be sufficient to offset additional losses of fixed nitrogen in a more productive Arctic is not known

2.2.3. Changes in the biological pump

The increase in the flux of CO2 from the atmosphere into the Arctic Ocean has tripled over the last 3 decades (from 24 to 66 Tg C yr-1) (Bates et al. 2006). This increase is attributed to the recent loss of sea ice that facilitated both increased primary production and sea-air CO2 exchange. The recent increase in primary production reported here should further enhance this exchange, due to the reduction in surface water pCO2 during the conversion of inorganic CO2 to organic carbon by phytoplankton that eventually sinks below the thermocline. Although it has been calculated that outgassing of CO2 will increase by 8 g C m-2 yr-1 for every 1ºC increase in sea surface temperature (Anderson and Kaltin 2001), the biologically-mediated decrease in surface pCO2 should partially offset the increased outgassing of CO2 expected as Arctic surface waters warm in upcoming years. In fact, it has been suggested that when anticipated changes in CO2 solubility (due to changes in both temperature and salinity) and phytoplankton production are taken into account, the potential for the Arctic Ocean to act as a sink for atmospheric CO2 will increase in the future (Anderson and Kaltin 2001). However, longer-term observations are required to understand the extent to which primary production will be either intensified or weakened by the many concurrent environmental changes ongoing in the Arctic Ocean (e.g., declines in sea ice cover, increased SST, increased freshwater fluxes, changes in nutrient and light availability).

2.3. Changes in Arctic ozone and UV radiation

In recent years, stratospheric ozone levels over the Arctic have been diminishing, although not to the extent as that observed in the Antarctic (Solomon et al. 2007). As a result, the surface of the Arctic Ocean is being exposed to increasingly high levels of UV radiation (UVR), which will impact surface biological communities as well as rates of photooxidation of CDOM. The extent to which these processes will intensify as sea ice cover diminishes is currently unknown, but in serious need of increased attention.

2.3.1. Impacts on biological community

Due to their high dissolved oxygen levels and intense UVR during the summer, Arctic surface waters are strongly oxidizing ecosystem (Camus and Gulliksen 2005). Incoming UVR levels are controlled by solar zenith angle, cloudiness and stratospheric ozone concentration. The relative amount of UVR to visible light increases with heavy cloud cover (Eilertsen and Holm-Hansen 2000). Changes in CDOM influence the penetration of UV radiation into the surface ocean, with major consequences for aquatic biogeochemical processes (Epp et al. 2007, Retamal et al. 2008). The deepest penetration of UVB into the water column of the Kongsfjord has been determined at 10 m depth (Bischof 2000). Optical measurements and modeling results indicate that ozone-related UVR influences on food web processes in the Arctic Ocean are likely to be small relative to the effects caused by variation in the concentrations of CDOM (Gibson et al. 2000).

In shallow water regions, macrophytes are often exposed to high doses of UVR during the summer months. However, large differences in UVR sensitivity have been observed for a variety of physiological parameters, with lower subtidal species being more sensitive to enhanced UVR treatments than upper subtidal species (Bischof 2000, Karsten et al. 2001, Van de Poll et al. 2002, Holzinger et al. 2004). Only minor differences were found between Arctic and temperate isolates, suggesting that specific differences in UVR sensitivity have not evolved in Arctic species (Van de Poll et al. 2002). Photosynthetic electron transport rates are particularly sensitive to UVR exposure, which is partly a result of the degradation of the CO2-fixing Rubisco protein (Bischof 2000). Decreased photochemical efficiency of photosystem II have also been observed under enhanced UVR (Holzinger et al. 2004). Microscopic examinations show that the photosynthetic apparatus can be severely damaged by UVR, including wrinkling of thylakoid membranes, lumen dilatations, mitochondrial damage, and alterations of outer membranes (Holzinger et al. 2004). However, after some days of treatment, macroalgae are able to recover and UVR inhibition is reduced (Bischof 2000).

Less is known about the impacts of increased UVR of Arctic phytoplankton. Photosynthesis in near-surface waters of the Barents Sea, determined from incubation experiments, has been shown to be inhibited by 50-75% due to UVR, with detectable inhibition down to a depth of 10 m (Eilertsen and Holm-Hansen 2000). The haptophyte Phaeocystis pouchetii, a dominant member of the phytoplankton community, was less sensitive to UVR than centric diatoms. Although the calculated critical depth required to produce an increase in phytoplankton biomass was shallower during periods of elevated UVR, total annual production was not severely affected. This is consistent with a model of UVR inhibition of primary production for western Spitsbergen which showed that, on average, water column primary production was reduced by only 2.9% because of to enhanced UVB (Wangberg et al. 2006). Even if stratospheric ozone concentrations were reduced to 200 DU, UVR inhibition would increase to just 4.4%.

Studies of the impacts of UVR on Arctic zooplankton have focused largely on crustaceans and their response to UVR exposure and increased oxygen radical production. Camus and Gulliksen (2005) investigated three amphipod species (one from the deep-sea, one from the benthic sublittoral zone, and one from the ice pack) and found that the deeper species were much more susceptible to oxidative stress. These species exhibited reduced oxygen radical scavenging capacity while the pack ice amphipod (Gammarus wilkitzkii) possessed a mechanism that prevents the diffusion of exogenous reactive oxygen through the gills and allows excretion of internal hydrogen peroxide through the gills into the environment (Camus and Gulliksen 2005).

In addition, carnivorous and herbivorous amphipods exhibit different responses to UVR stress. Carnivorous amphipods exhibited elevated sensitivity to UVR exposure which was accompanied by degradation of tissue carotenoid and mycosporine-like amino acids (MAA), compounds used as UV photoprotective sunscreens, and a decrease of the enzymatic antioxidant defenses (Obermueller et al. 2005). In contrast, herbivorous amphipods were protected by high concentrations of MAAs obtained from their algal diet, and no oxidative stress occurred under experimental UVR. The species-specific degree of UV tolerance correlates well with the vertical distribution in the water column (Obermueller et al. 2005).

2.3.2. Impacts of UV photodegradation of CDOM

Waters of the Arctic Ocean are characterized by high concentrations CDOM. The dissolved organic matter (DOM) that enters the ocean either through terrestrial runoff or in situ production by phytoplankton impacts the optical properties of coastal seawater (Matsuoka et al. 2007) and affects carbon cycling on a global scale (Moran et al. 2000). UV-B radiation has been shown to strongly influence degradation of this material, impacting the biogeochemical cycling of aquatic carbon, nitrogen, sulfur, as well as metals important to a wide range of life processes. By changing the biological availability of DOM to microorganisms, UV-B radiation accelerates its photochemical transformation into dissolved inorganic carbon (DIC) and nitrogen, including carbon monoxide (CO), carbon dioxide, and ammonium (Epp et al. 2007). Furthermore, the direct photooxidation of CDOM by UVR may represent a significant source of DIC in the ocean (Johannessen et al. 2001).

Water temperature, as well as the origin and light history of CDOM, strongly affect the efficiency of CO photoproduction. Terrestrial CDOM is more efficient at photochemically producing CO than is phytoplankton-derived CDOM (Zhang et al. 2006). However, Johannessen et al. (2001) noted that the efficiency of DIC production increases offshore, which suggests that the most highly absorbing and quickly faded terrestrial chromophores are not those directly responsible for DIC photoproduction. Biological degradation of photobleached DOM is more rapid than that of unbleached material, and this net positive effect was evident even for extensively photodegraded material (Moran et al. 2000). These factors should be taken into account in modeling the photochemical fluxes of CO and other related CDOM photoproducts on varying spatiotemporal scales (Zhang et al. 2006).

2.4. Quantification of air-Sea CO2 exchange

Air-sea exchange of CO2 is generally calculated from measurements of the difference in the partial pressure of CO2 (pCO2) between the ocean and atmosphere (pCO2), CO2 solubility in seawater, and surface wind speed. Data from Arctic waters are sparse, but indicate that this region is a net sink for atmospheric CO2 due to cooling of surface waters and areas of high biological productivity.

Data from the shelf and slope waters of the western Arctic measured during the summers of 1998-2000 suggest that this continental margin is a moderate to strong sink for atmospheric CO2 in the summer. However, the contribution of biological production to the low surface water pCO2 was small, with physical factors such as water mixing and cooling dominating air-sea CO2 exchange (Murata et al. 2003). Similarly, the Barents Sea was a net sink of atmospheric CO2 (Fransson et al. 2001, Katlin et al. 2002, Omar et al. 2007) with cooling of surface waters being the primary factor responsible driving the air-sea flux (Katlin et al. 2002). In the Greenland Sea and the Barents Sea, the annual net ocean CO2 sink was estimated to be 5220 g C m-2 yr-1 and 4618 g C m-2 yr-1, respectively, reaching its maximum in winter and minimum in summer. In this region, the wind speed and pCO2 exerted a greater influence on the seasonal variation than the sea ice coverage, with the air sea flux being positively correlated to the North Atlantic Oscillation (Nakaoka et al. 2006). Nitishinsky et al. (2007) reported that the Laptev Sea Shelf takes up 25.2 mg CO2 m-2 d-1 from atmosphere, whereas the western part of the East-Siberian Sea Shelf loose 3.6 mg CO2 m-2 d-1 to the atmosphere. These spatial differences were attributed to variations in bottom topography, river runoff, exchange with surrounding seas, and the wind field.

In contrast to those Arctic regions where the air-sea CO2 flux is dominated by physical processes, Miller et al. (2002) observed a 250 µmol kg-1 drop in surface DIC concentrations in the North Water polynya between winter and summer, which they attributed to biological drawdown of CO2. They concluded that although the surface waters were supersaturated with CO2 in early spring, extensive ice cover limited the outgassing of CO2. In early autumn, when ice again was beginning to cover the area, the surface waters were still undersaturated in CO2, implying that the North Water polynya is a net sink of atmospheric carbon. Similar observations have been reported for the Chukchi Sea (Bates 2006) and Canada Basin (Bates et al. 2006), where sea ice cover restricts air-sea gas exchange in the autumn, winter, and early spring, and high rates of primary productivity in the summer reduces pCO2 to 80-220 µatm. It has been estimated that the Arctic Ocean sink for CO2 has tripled over the last 3 decades due to substantial sea-ice retreat during that time (Bates et al. 2006).

Highly productive regions of the Arctic Ocean, such as the Chukchi Sea, are likely to remain perennial sinks for atmospheric CO2. However, as rates of primary productivity increase in other regions of the Arctic Ocean due to declining sea ice cover, their status as an atmospheric CO2 sink remains to be determined.

2.5. Ocean acidification

Studies of ongoing and anticipated increases in surface ocean pH (ocean acidification) have focused primarily on warm, low latitude regions where reef-building corals are common. However, a few studies also have looked at the impact of pH changes in high latitude waters where calcifying coccolithophorids and planktonic pteropods are important members of the surface ocean community. Recently, model results from Orr et al. (2005) suggested that within the next 50 years, some polar and subpolar surface waters will become undersaturated with respect to carbonate at atmospheric CO2 levels much lower than had been previously thought (approximately 750 ppm). Furthermore, they calculated that in parts of the subarctic Pacific, the aragonite saturation horizon could shoal from 120 m to the surface, while in the North Atlantic (north of 50°N), it could shoal from 2,600 m to only 115 m. These large changes in the North Atlantic are due to deeper penetration and higher concentrations of anthropogenic CO2, a trend supported by recent data- and model-based estimates (Sarmiento et al., 1992, Gruber et al. 1998, Orr et al., 20001, Sabine et al. 2004). These large changes in future carbonate saturation should have large, but as yet unexplored, impacts on calcifying organisms in Arctic waters.

For example, during the late 1990’s and early 21st century, coccolithophorid blooms dominated the Bering Sea phytoplankton community for much of the summer (Stockwell et al. 2001, Lida et al. 2002, Merico et al. 2003). These unusual observations were thought to be the result of anomalous physical conditions, including strong spring storms followed by a calm summer, shallow mixed layers, and enhanced freshwater content (Stabeno et al. 2001). Because calcifying organisms have a large impact on the alkalinity and the upper ocean carbon cycle, whether these blooms are likely to recur in the future is an important question.

Early experiments showed that the coccolithophores Emiliana huxleyi and Gephyrocapsa oceanica exhibited reduced rates of calcification and malformed coccoliths when grown at reduced pH levels intended to simulate increased atmospheric CO2 (Riebesell 2000). These results suggested that the prevalence of these calcifying organisms was likely to diminish in the future as atmospheric CO2 increases and the pH of the surface ocean drops. However, more recent research suggests just the opposite (Iglesias-Rodriguez 2008). In these experiments, the pH was manipulated by increasing the CO2 concentration directly (Riebesell et al. 2000 added HCl to decrease the pH). As a result, E. huxleyi exhibited increased rates of calcification and net primary production, which was attributed to the higher CO2 partial pressures. Given these conflicting results, it is still not clear what impact ocean acidification is likely to have on the calcifying organisms that dominate the Arctic and subarctic and how this will impact the cycling of carbon.

2.6. Validation of remote sensing and modeling products

Because of its geographic location, the Arctic Ocean is a particularly challenging region for retrieval of satellite-derived information from its surface waters. Large solar zenith angles make atmospheric correction challenging and heavy cloud cover makes direct comparisons between satellite and in situ observations difficult. Furthermore, because of its relatively high CDOM content, retrievals of Chl a using standard empirical algorithms pose particularly difficult problems. Semi-analytical algorithms that provide independent estimates of CDOM, Chl a, and backscatter may prove to be more reliable (e.g. Maritorena et al. 2002), but these have yet to validated for Arctic waters.

In addition, NASA supports the development and implementation of biogeochemical models that can be used to understand and predict changes in the major elemental cycles (particularly carbon) that are likely to be manifested in the face of continued anthropogenic perturbations. Like satellite data products, these model results must be rigorously validated using data collected in situ.

3. References

Aagaard, K., L K. Coachman, and E. Carmack (1981), On the halocline of the Arctic Ocean, Deep-Sea Res., Part A, 28, 529-545.
Anderson, L. G., and S. Kaltin (2001), Carbon fluxes in the Arctic Ocean - potential impact by climate change, Polar Res., 20(2), 225-232.
Arctic Climate Impact Assessment (ACIA) (2005), Cambridge University Press, Cambridge, United Kingdom and New York, NY, USA.
Arrigo, K. R., G. L. van Dijken, and S. Bushinsky (2008), Primary Production in the Southern Ocean, 1997-2006. J. Geophys. Res.. 113, doi:10.1029/2007JC004551.
Bates, N. R. (2006), Air-sea CO2 fluxes and the continental shelf pump of carbon in the Chukchi Sea adjacent to the Arctic Ocean. J. Geophys. Res. 111(C10), C1001.
Bates, N. R., S. B. Moran, D. A. Hansell, and J. T. Mathis (2006), An increasing CO2 sink in the Arctic Ocean due to sea ice loss, Geophys. Res. Lett., 33(23), L23609, doi:10.1029/2006GL027028.
Bischof, K. (2000), Effects of enhanced UV-radiation on photosynthesis of Arctic/cold-temperate macroalgae. Berichte zur Polarforschung, 375, i-xiii, 1-88.
Camus, L. and B. Gulliksen (2005), Antioxidant defense properties of Arctic amphipods: comparison between deep-, sublittoral and surface-water species. Mar. Biol. 146(2), 355-362.
Carmack, E. C. and D. Chapman (2003), Wind-driven shelf/basin exchange on an Arctic shelf: The joint roles of ice cover extent and shelf-break bathymetry, Geophys. Res. Lett., 30, 1778, doi:10.1029/2003GL017526.
Carmack, E. C., and R. W. Macdonald (2002), Oceanography of the Canadian shelf of the Beaufort Sea: A setting for marine life, Arctic, 55, 29-45.
Carmack, E. C., K. Aagaard, J. H. Swift, R. W. MacDonald, F. A. McLaughlin, E. P. Jones, R. G. Perkin, J. N. Smith, K. M. Ellis, and L. R. Killius (1997), Changes in temperature and tracer distributions within the Arctic Ocean: results from the 1994 Arctic Ocean Section, Deep-Sea Res., Part II, 44(8), 1487-1502.
Chapman, W. L., and J. E. Walsh (2003), Observed climate change in the Arctic, updated from Chapman and Walsh, 1993: Recent variations of sea ice and air temperatures in high latitudes, Bull. Am. Met. Soc., 74(1), 33-47.
Codispoti, L. A. (1979), Arctic Ocean processes in relation to the dissolved silicon content of the Atlantic, Mar. Sci. Commun., 5(6), 361-381.
Comiso, J. C. (2006), Abrupt decline in Arctic winter sea ice cover, Geophys. Res. Lett., 33, L18504.
Comiso J. C., C. L. Parkinson, R. Gersten, and L. Stock (2008), Accelerated decline in the Arctic sea ice cover, Geophys. Res. Lett., 35, L01703, doi:10.1029/2007GL031972.
Devol, A. H., L. A. Codispoti, and J.P. Christensen (1997), Summer and winter denitrification rates in western Arctic shelf sediments, Cont. Shelf Res., 17(9), 1029-1033.
Dickson, R. R., T. J. Osborn, J. W. Hurrell, J. Meincke, J. Blindheim, B. Adlandsvik, T. Vinje, G. Alekseev, and W. Maslowski (2000), Arctic ocean response to the North Atlantic oscillation, J. Clim., 13, 2671-2696.
Eilertsen, H. C. and O. Holm-Hansen (2000), Effects of high latitude UV radiation on phytoplankton and nekton modelled from field measurements by simple algorithms. Polar Res. 19(2), 173-182.
Epp, R. G., D. J. Erickson, N. D. Paul, and B. Sulzberger (2007), Interactive effects of solar UV radiation and climate change on biogeochemical cycling. Photochem. Photobiol. Sci., 6(3), 286-300.
Falk-Petersen, S., H. Hop, W. Budgell, E. Hegseth, R. Korsnes, T. Loyning, J. Orbaek, T. Kawamura, and K. Shirasawa (2000), Physical and ecological processes in the marginal ice zone of the northern Barents Sea during the summer melt period, J. Mar. Sys., 27(1-3), 131-159.
Fransson, A., M. Chierici, L. C. Anderson, I. Bussmann, G. Kattner, E. P. Jones, and J. H. Swift (2001), The importance of shelf processes for the modification of chemical constituents in the waters of the Eurasian Arctic Ocean: implication for carbon fluxes. Cont. Shelf Res. 21(3), 225-242.
Gibson, J. A.E., B. Nieke, R. Pienitz, W. F. Vincent (2000), Control of biological exposure to UV radiation in the Arctic Ocean: comparison of the roles of ozone and riverine dissolved organic matter. Arctic 3(4), 372-382.
Gordeeva, V. V. (2006), Fluvial sediment flux to the Arctic Ocean. Geomorphology, 80 (1-2), 94-104.
Gosselin, M., M. Levasseur, P. A. Wheeler, R. A. Horner, and B. C. Booth (1997), New measurements of phytoplankton and ice algal production in the Arctic Ocean, Deep-Sea Res., Part II, 44(8), 1623-1644.
Gruber, N. (1998), Anthropogenic CO2 in the Atlantic Ocean. Glob. Biogeochem. Cycles 12, 165–191.
Harrison, W. and G. Cota (1991), Primary production in polar waters - relation to nutrient availability, Polar Res., 10(1), 87-104.
Hill, V. J. (2008), The impacts of chromophoric dissolved organic material on surface ocean heating in the Chukchi Sea, J. Geophys. Res., doi:10.1029/2007JC004119.
Hill, V. and G. Cota (2005), Spatial patterns of primary production on the shelf, slope and basin of the Western Arctic in 2002, Deep-Sea Res., Part II, 52(24), 3344-54.
Holmes, R. M., J. W. McClelland, P. A. Raymond, B. B. Frazer, B. J. Peterson, and M. Stieglitz (2008), Lability of DOC transported by Alaskan rivers to the Arctic Ocean. Geophys. Res. Lett., 35(3) L03402, doi:10.1029/2007GL032837.
Holzinger, A., C. Lutz, U. Karsten, and C. Wiencke (2004), The effect of ultraviolet radiation on ultrastructure and photosynthesis in the red macroalgae Palmaria palmata and Odonthalia dentata from Arctic waters. Plant Biol. 6(5), 568-577.
Huntington, H. P. and S. Fox (2005), The Changing Arctic: Indigenous Perspectives, Chapter 3 in Arctic Climate Impact Assessment (ACIA), pp. 61-98, Cambridge University Press, New York.
Iglesias-Rodriguez, M. D., P. R. Halloran, R. E. M. Rickaby, I. R. Hall, E. Colmenero-Hidalgo, J. R. Gittins, D. R. H. Green, T. Tyrrell, S. J. Gibbs, P. von Dassow, E. Rehm, E. V. Armbrust, and K. P. Boessenkool (2008), Phytoplankton calcification in a high-CO2 world. Science 320(5874), 336-340.
Jakobsson, M., A. Grantz, Y. Kristoffersen, and R. Macnab (2003), Physiographic provinces of the Arctic Ocean seafloor, Bull. Geol. Soc. Am., 115(12), 1443-1455.
Johannessen, S.C. and W. L. Miller (2001), Quantum yield for the photochemical production of dissolved inorganic carbon in seawater. Mar. Chem., 76(4), 271-283.
Kaltin, S., L. G. Anderson, K. Olsson, A. Fransson, and M. Chierici (2002), Uptake of atmospheric carbon dioxide in the Barents Sea. J. Mar. Sys. 38(1-2), 31-45.
Karsten, U., K. Bischof, and C. Wiencke (2001), Photosynthetic performance of Arctic macroalgae after transplantation from deep to shallow waters. Oecologia 127(1), 11-20.
Legendre, L., S. F. Ackley, G. S. Dieckmann, B. Gullicksen, R. Horner, T. Hoshiai, I. A. Melnikov, W. S. Reeburgh, M. Spindler and C. W. Sullivan (1992), Ecology of sea ice biota: 2. Global significance. Polar Biol., 12, 429-444.
Levi, B. G. (2000), The decreasing Arctic ice cover, Physics Today, 53(1), 19-20.
Lindsay, R. W. and J. Zhang (2005), The thinning of Arctic sea ice, 1988-2003: Have we passed a tipping point? J. Clim., 18, 4879-4894.
Macdonald, R. (1996), Oceanography - Awakenings in the Arctic, Nature, 380(6572), 286-287.
Macdonald, R. W., S. M. Solomon, R. E. Cranston, H. E. Welch, M. B. Yunker, and C. Gobeil (1998), A sediment and organic carbon budget for the Canadian Beaufort shelf, Mar. Geol., 144(4), 255-273.
Maritorena S., D.A. Siegel and A. Peterson (2002), Optimization of a Semi-Analytical Ocean Color Model for Global Scale Applications. Appl. Opt. 41(15), 2705-2714.
Maslowski, W., D. C. Marble, W. Walczowski, and A. J. Semtner (2001), On large-scale shifts in the Arctic Ocean and sea-ice conditions during 1979-98, Ann. Glaciol., 33, 545-550.
Maslowski, W., D. Marble, W. Walczowski, U. Schauer, J. L. Clement, and A. J. Semtner (2004), On climatological mass, heat, and salt transports through the Barents Sea and Fram Strait from a pan-Arctic coupled ice-ocean model simulation, J. Geophys. Res., 109(C3), C03032, doi:10.1029/2001JC001039.
Matsuoka, A., Y Huot, K. Shimada, S-I. Saitoh, and M. Babin, (2007), Bio-optical characteristics of the western Arctic Ocean: Implications for ocean color algorithms. Can. J. Remote Sens., 33(6), 503-518.
Maslanik, J., S. Drobot, C. Fowler, W. Emery, and R. Barry (2007), On the Arctic climate paradox and the continuing role of atmospheric circulation in affecting sea ice conditions, Geophys. Res. Lett., 34, L03711.
Menard, H. W. and S. M. Smith (1966), Hypsometry of ocean basin provinces, J. Geophys. Res., 71(18), 4305-4325.
Merico, A., T. Tyrrell, C. W. Brown, S. B. Groom, and P. I. Miller (2003), Analysis of satellite imagery for Emiliania huxleyi blooms in the Bering Sea before 1997. Geophys. Res. Lett., 30(6), 1337.
Miller, L. A., P. L. Yager, K. A. Erickson, D. Amiel, .J Bacle, J. K. Cochran, M. E. Garneau, M. Gosselin, D. J. Hirschberg, B. Klein, B. LeBlanc, and W. L. Miller (2002), Carbon distributions and fluxes in the North Water, 1998 and 1999. Deep-Sea Res., Part II, 49(22-23), 5151-70.
Mock, T. and R. Gradinger (1999), Determination of Arctic ice algal production with a new in situ incubation technique, Mar. Ecol. Prog. Ser., 177, 15-26.
Moran, M. A., W. M. Sheldon, and R. G. Zepp, (2000), Carbon loss and optical property changes during long-term photochemical and biological degradation of estuarine dissolved organic matter. Limnol. Oceanogr., 45(6), 1254-1264.
Morehead, M. D., J. P. Syvitski, E. W. Hutton, and S. D. Peckham, (2003), Modeling the temporal variability in the flux of sediment from ungauged river basins. Glob. Planet. Change. 39, 95–110.
Murata, A., and T. Takizawa (2003), Summertime CO2 sinks in shelf and slope waters of the western Arctic Ocean, Cont. Shelf Res., 23(8), 753-776.
Nakaoka, S., S. Aiki, T. Nakazawa, G. Hashida, S. Morimoto, T. Yamanouchi, and H. Yoshikawa-Inoue (2006), Temporal and spatial variations of oceanic pCO2 and air-sea CO2 flux in the Greenland Sea and the Barents Sea. Tellus Series B Chemical and Physical Meteorology; 58(2), 148-161.
Napp, J. M. and G. L. Hunt, (2001), Anomalous conditions in the south-eastern Bering Sea 1997: Linkages among climate, weather, ocean, and biology. Fish. Oceanogr., 10(1), 61-68.
National Snow and Ice Data Center (NSIDC) (2007), Arctic sea ice shatters all previous record lows. http://nsidc.org/news/press/2007_seaiceminimum/20071001_pressrelease.html.
Niebauer, H., A. V, and S. Henrichs (1990), Physical and biological oceanographic interaction in the spring bloom at the Bering Sea marginal ice edge zone, J. Geophys. Res., 95(C12), 22229-22241.
Nitishinsky, M., L. G. Anderson, and J. A. Holemann, (2007), Inorganic carbon and nutrient fluxes on the Arctic Shelf. Cont. Shelf Res., 27(10-11), 1584-1599.
Obermueller, B., U. Karsten, and D. Abele (2005), Response of oxidative stress parameters and sunscreening compounds in Arctic amphipods during experimental exposure to maximal natural UVB radiation. J. Exp. Mar. Biol. Ecol. 323(2), 100-117.
Omar, A. M., T. Johannessen, A. Olsen, S. Kaltin, and F. Rey (2007), Seasonal and interannual variability of the air-sea CO2 flux in the Atlantic sector of the Barents Sea. Mar. Chem. 104(3-4), 203-213.
Orr, J. C., E. Maier-Reimer, U. Mikolajewicz, P. Monfray, J. L. Sarmiento, J. R. Toggweiler, N. K. Taylor, J. Palmer, N. Gruber, C. L. Sabine, C. Le Quere, R. M. Key, and J. Boutin, (2001), Estimates of anthropogenic carbon uptake from four three-dimensional global ocean models. Glob. Biogeochem. Cycles 15, 43–60.
Orr, J. C., V. J. Fabry, O. Aumont, L. Bopp, S. C. Doney, R. A. Feely, A. Gnanadesikan, N. Gruber, A. Ishida, F. Joos, R. M. Key, K. Lindsay, E. Maier-Reimer, R. Matear, P. Monfray, A. Mouchet, R. G. Najjar, G.-K. Plattner, K. B. Rodgers, C. L. Sabine, J. L. Sarmiento, R. Schlitzer, R. D. Slater, I. J. Totterdell, M.-F. Weirig, Y. Yamanaka and A. Yool (2005), Anthropogenic ocean acidification over the twenty-first century and its impact on calcifying organisms. Nature 437, 681-686.
Overland, J. E. and M. Wang (2007), Future regional Arctic sea ice declines, Geophys. Res. Lett., 34, L17705.
Pabi, S., G. L. van Dijken, and K. R. Arrigo (2008), Primary Production in the Arctic Ocean, 1998-2006. J. Geophys. Res. 113. doi:10.1029/2007JC004578.
Parkinson, C. L. (2000), Variability of Arctic sea ice: The view from space, and 18-year record. Arctic, 53, 341-358.
Perovich, D. K. B. Light, H. Eicken, K. F. Jones, K. Runciman, and S. V. Nghiem (2007), Increasing solar heating of the Arctic Ocean and adjacent seas, 1979-2005: Attribution and role in the ice-albedo feedback, Geophys. Res. Lett., 34, L19505.
Peterson, B. J., J. McClelland, R. Curry, R. M. Holmes, J. E. Walsh, and K. Aagaard (2006), Trajectory shifts in the Arctic and subarctic freshwater cycle, Science, 313, 1061-1066.
Rahmstorf, S. (2002), Ocean circulation and climate during the past 120,000 years. Nature 419, 207-214.
Retamal, L., W. F. Vincent, C. Martineau, and C. L. Osbum (2007 ), Comparison of the optical properties of dissolved organic matter in two river-influenced coastal regions of the Canadian Arctic. Estuarine, Coast. Shelf Sci., 72(1-2), 261-272.
Retamal, L., S Bonilla, and W. F. Vincent (2008), Optical gradients and phytoplankton production in the Mackenzie River and the coastal Beaufort Sea. Polar Biol. 31(3), 363-379.
Riebesell, U., I. Zondervan, B. Rost, P. D. Tortell, R. E. Zeebe, and F. M. M. Morel (2000), Reduced calcification of marine plankton in response to increased atmospheric CO2. Nature, 407(6802), 364-367.
Rigor, I. G. and J. M. Wallace (2004), Variations in the age of Arctic sea-ice and summer sea-ice extent, Geophys. Res. Lett., 31, L09401.
Rigor, I. G., J. M. Wallace, and R. L. Colony (2002), Response of sea ice to the Arctic Oscillation, J. Climate, 15(18), 2648.
Rothrock, D. A. and J. Zhang (2005), Arctic Ocean sea ice volume: What explains its recent depletion? J. Geophys. Res., 110, C01002.
Rothrock, D. A., Y. Yu, and G. A. Maykut (1999), Thinning of the Arctic sea ice cover, Geophys. Res. Lett., 26(23), 3469.
Sabine, C. L., R. A. Feely, N. Gruber, R. M. Key, K. Lee, J. L. Bullister, R. Wanninkhof, C. S. Wong, D. W. R. Wallace, B. Tilbrook, F. J. Millero, T.-H. Peng, A. Kozyr, T. Ono, and A. F. Rios (2004),The ocean sink for anthropogenic CO2. Science 305, 367–370.
Sakshaug, E., K. Andresen, S. Myklestad, and Y. Olsen (1983), Nutrient status of phytoplankton communities in Norwegian waters (marine, brackish and fresh) as revealed by their chemical composition, J. Plank. Res., 5(2), 175-196.
Sakshaug, E. (2003), Primary and secondary production in the Arctic Seas, in The organic carbon cycle in the Arctic Ocean.edited by Stein R, Macdonald RW, pp. 57-81, Springer-Verlag, Berlin.
Sarmiento, J. L., J. C. Orr, and U. Siegenthaler, (1992), A perturbation simulation of CO2 uptake in an ocean general circulation model. J. Geophys. Res., 97, 3621–3645
Serreze, M. C., M. M. Holland, and J. Stroeve (2007), Perspectives on the Arctic’s shrinking sea-ice cover, Science, 315, 1533-1536 (2007).
Shiklomanov, I. A. (2000), Appraisal and assessment of world water resources, Water International, 25(1), 11.
Shimada, K., T. Kamoshida, M. Itoh, S. Nishino, E. Carmack, F. McLaughlin, S. Zimmerman, and A. Proshutinsky (2006), Pacific Ocean inflow: Influence on catastrophic reduction of sea ice cover in the Arctic Ocean, Geophys. Res. Lett., 33, L08605.
Smith, D. M. (1998), Recent increase in the length of the melt season of perennial Arctic sea ice, Geophys. Res. Lett., 25(5), 655-8.
Solomon, S., R. W. Portmann, and D. W. J. Thompson (2007), Contrasts between antarctic and arctic ozone depletion. Proc. Nat. Acad. Sci. 104(2), 445-449.
Stabeno, P. J., N. A. Bond, N. B. Kachel, S. A. Salo, and J. D. Schumacher (2001), On the temporal variability of the physical environment over the south-eastern Bering Sea. Fish. Oceanogr., 10(1), 81-98.
Steele, M. and T. J. Boyd (1998), Retreat of the cold halocline layer in the Arctic Ocean, J. Geophys. Res., 103, 10419-10435.
Stockwell, D. A., T. E. Whitledge, S. I. Zeeman, K. O. Coyle, J. M. Napp, R. D. Brodeur, A. I. Pinchuk, and G. L. Hunt, (2001), Anomalous conditions in the south-eastern Bering Sea, 1997: Nutrients, phytoplankton and zooplankton. Fish. Oceanogr., 10(1), 99-116.
Subba Rao, D. V. and T. Platt (1984), Primary productivity of Arctic waters, Polar Biol., 3, 191-201.
Syvitski, J. P. M. (2002) Sediment discharge variability in Arctic rivers: implications for a warmer future. Polar Res., 21(2), 323-330.
Tanaka, T., L. D. Guo, C. Deal, N. Tanaka, T. Whitledge, and A. Murata (2004), N deficiency in a well-oxygenated cold bottom water over the Bering Sea shelf: influence of sedimentary denitrification, Cont. Shelf Res., 24(12), 1271-1283.
Tyrrell, T. (1999), The relative influences of nitrogen and phosphorus on oceanic primary production, Nature, 400(6744), 525-538.
Van Blarcum, S., J. R. Miller, and G. L. Russell (1995), High latitude river runoff in a doubled CO2 climate. Clim. Change. 30(1), 7-26.
Van de Poll, W. H., A. Eggert, A. G. J. Buma, and A. M. Breeman (2002), Temperature dependence of UV radiation effects in Arctic and temperate isolates of three red macrophytes. Euro. J. Phycol. 37(1), 59-68.
Wangberg, SA, K. I. M. Andreasson, K. Garde, K. Gustavson, P. Henriksen, and T. Reinthaler (2006), Inhibition of primary production by UV-B radiation in an arctic bay - model calculations. Aquatic Sci. 68(2), 117-128.
Yamamoto-Kawai, M., E. Carmack, and F. McLaughlin (2006), Nitrogen balance and Arctic throughflow, Nature, 443, 43.
Zehr, J. P., M. Wyman, V. Miller, L. Dugay, and D. G. Capone (1993), Modification of the iron protein of nitrogenase in natural populations of Trichodesmium thiebautii, App. Env. Microbiol., 59(3), 669-676.
Zhang, Y, H. X. Xie, and G. H. Chen (2006), Factors affecting the efficiency of carbon monoxide photoproduction in the St. Lawrence estuarine system (Canada). Env. Sci. Tech., 40(24), 7771-7777.