Carbon cycle and climate change in the Arctic

May 7, 2012, 11:32 am

This is Section 9.5 of the Arctic Climate Impact Assessment
Lead Author: Harald Loeng; Contributing Authors: Keith Brander, Eddy Carmack, Stanislav Denisenko, Ken Drinkwater, Bogi Hansen, Kit Kovacs, Pat Livingston, Fiona McLaughlin, Egil Sakshaug;  Consulting Authors: Richard Bellerby, Howard Browman,Tore Furevik, Jacqueline M. Grebmeier, Eystein Jansen, Steingrimur Jónsson, Lis Lindal Jørgensen, Svend-Aage Malmberg, Svein Østerhus, Geir Ottersen, Koji Shimada

 

The Arctic Ocean has not been considered a significant carbon sink; first, because extensive sea-ice cover constrains atmosphere–ocean exchange, and second, because levels of biological production under perennial sea ice were considered low[1]. Under warmer conditions, however, the amount of carbon sequestered by the Arctic Ocean is very likely to increase significantly. The role of the Arctic as a potential carbon source, in the form of CH4 and CO2, is unclear owing to limited information on the likely impact of climate change on the substantial frozen reserves in permafrost and gas hydrate layers.

The ocean carbon cycle comprises a physical pump, a biological pump, and an alkalinity or anion pump. The physical pump is driven by physical and chemical processes, which affect the solubility of CO2 and the transport of water from the surface mixed layer to depth. The biological pump is driven by primary production, consuming dissolved CO2 through photosynthesis and producing particulate organic carbon (POC) and DOC. The alkalinity pump concerns the removal of carbon by calcification in the upper waters and the release of carbon when calcium carbonate is dissolved at depth. The alkalinity pump is not affected by temperature itself, but is affected indirectly through shifts in biological speciation.

Physical pump (9.5.1)

The presence of sea ice strongly affects the physical pump, which regulates the exchange of CO2 between the atmosphere and the ocean. This exchange is primarily determined by the difference in partial pressure of CO2 (pCO2) over the air–sea interface. Physical factors, such as wind mixing, temperature, and salinity, are also important in this exchange. Dissolved inorganic carbon (DIC) is the largest component of the marine carbon pool.

Multi-year ice restricts air–sea exchange over the central Arctic Ocean and seasonal sea ice restricts air–sea exchange over shelf regions to ice-free periods. Because the solubility of CO2 in seawater increases with decreasing temperature, the largest uptake of atmospheric CO2 occurs primarily in the ice-free Nordic Seas (~86 x 1012 g C/yr[2]) where northward flowing Atlantic waters are rapidly cooled. Similarly, the Barents and Bering/Chukchi Seas, where inflowing Atlantic and Pacific waters undergo cooling, are also important uptake regions: uptake in the Barents Sea is ~9 x 1012 g C/yr[3] and in the Bering/Chukchi Seas is ~22 x 1012 g C/yr[4]. Uptake in the Bering/Chukchi Seas is higher than in the Barents Sea for reasons discussed in greater detail in section 9.5.2; namely, a higher potential for new production owing to a greater supply of nutrients, and a larger area of retreating ice edge along which much of the primary production occurs. Carbon uptake in the ice-covered Arctic Ocean and interior shelf seas is ~31 x 1012 g C/yr[5].

Although these fluxes are not large on a global scale (~2,000 x 1012 g C/yr), the air–sea CO2 flux is very likely to increase regionally under scenarios of climate warming. For example, the ACIA-designated models project the Barents Sea and the northern Bering Sea to be totally ice-free by 2050 (see section 9.2.5.2). Such changes in ice cover and longer periods of open water will result in more regions that resemble the Greenland Sea, where the physical pump is strong due to low surface water temperatures and high wind speeds[6]. Atmospheric exchange will also increase as the areal coverage of the permanent ice pack is reduced and more leads and polynyas are formed. Here, the combination of increased atmospheric exchange (driven by winds) and ventilation (driven by sea-ice formation and convection) transport CO2 from the atmosphere into the halocline and potentially deeper, eventually entering the deep North Atlantic Ocean and the THC. Ventilation of Arctic Ocean intermediate waters has been estimated to sequester ~0.026 Gt C/yr, nearly an order of magnitude more than the sink due to convection in the Greenland Sea[7] and this is very likely to increase, possibly significantly.

Seasonally ice-covered shelf regions are also important dense water formation areas. Brine release during sea-ice formation increases the density of surface waters which then sink and are advected from the shelf to basin interiors, transporting CO2 into the halocline and deeper waters. Under warming conditions, ice formation on shelves will occur later and ice melt will occur earlier, thereby increasing the time available for air–sea interaction/equilibration and CO2 uptake. The coincidence of open water with late summer storms will also increase air–sea exchange and CO2 uptake.

Changes in dense water production and the THC will affect the ocean carbon reservoir[8]. The global ocean stores approximately fifty times more carbon than the atmosphere, mostly in the deep waters of the Pacific Ocean owing to their volume and long residence time. Slowing or stopping the THC would make the Atlantic circulation more like that of the Pacific, increasing its carbon storage and thus weakening the greenhouse effect and cooling the atmosphere – a negative feedback. In contrast however, if sites of deep ventilation were to move northward into the Arctic Basin[9], the resulting overturn may result in a positive feedback due to CO2 release to the atmosphere.

Changes in ice cover extent also affect the uptake of atmospheric CO2 by altering the equilibrium concentrations in the water column. Anderson L. and Katlin[10], using the Roy et al.[11] solubility equations, calculated that melting 2 to 3 m of sea ice and mixing the resulting freshwater into the top 100 meters (m) of the water column would increase CO2 uptake and could remove ~3 g C/m2. But, where warming is sufficient to increase surface water temperatures by 1°C, ~8 g C/m2 could be released due to the decrease in solubility. At high latitudes, surface waters are often undersaturated because heat is lost to the atmosphere more quickly than CO2 can dissolve. If ice cover retreated and the contact period with the atmosphere increased, this undersaturation would result in atmospheric CO2 uptake. Anderson L. and Katlin[12], using data for the Eurasian Basin where Atlantic waters dominate the upper water column, calculated that surface waters in the St. Anna Trough, the Eurasian Basin, and the Makarov Shelf slope have a potential carbon uptake of 35, 48, and 7 g C/m2, respectively, when ice cover conditions allow saturation.

 

caption Fig. 9.34. Profiles of the fugacity (partial pressure corrected for the fact that the gas is not ideal) of CO2 in Canada Basin and the Eurasian Basin. Data to the left of the dotted line are undersaturated and to the right are over-saturated.

 

Regionally, the effects of upwelling of halocline waters onto the shelf must also be considered. For example, a profile of the fugacity (partial pressure corrected for the fact that the gas is not ideal) of CO2 (f CO2) shows that Pacific-origin waters below 50 m in Canada Basin are oversaturated due to their origin in the productive Bering/Chukchi Seas (see Fig. 9.34). If upwelling brought these oversaturated waters onto the shelf and they mixed with surface waters CO2 would be released. Upwelling of waters with salinity ~33 (near 150 m in the Canada Basin) has been observed on the Alaskan and Beaufort shelves[13]. Upwelling is also expected to increase when the ice edge retreats beyond the shelf break[14]. In contrast, the f CO2 profile of Atlantic-origin waters shows that waters below 50 m in the Eurasian Basin are undersaturated and will take up atmospheric CO2 if moved onto the shelf by upwelling[15]. Hence, the recent shift in the Makarov Basin from a Pacific- to an Atlantic-origin halocline has modified shelves on the perimeter from a potential source to a potential sink of atmospheric CO2.

Biological pump (9.5.2)

The DOC concentrations in the deep arctic regions are comparable to those in the rest of the world’s oceans[16]. Within the Arctic Ocean, shelves are regions of high biological production, especially those within the Bering, Chukchi, and Barents Seas. Here, CO2 uptake is increased because CO2 fixation during photosynthesis affects the physical pump by reducing pCO2.

Levels of primary production are high on shelves due to increased light levels during ice-free periods and the supply of new nutrients by advection or vertical mixing. Although phytoplankton blooms are patchy, they are strongly associated with the retreating ice edge and the position of the ice edge in relation to the shelf break. In the northern Bering and southern Chukchi Seas, primary production occurs over a shallow shelf (50 to 200 m) and as the zooplankton and bacterioplankton cannot fully deplete this carbon source, it is either transferred to the benthos or advected downstream[17]. On the southeast Bering Sea shelf, which is deeper at ~200 m, there is potential for a match/mismatch of primary production and zooplankton grazing due to water temperature. An early bloom in cold melt water means most of the primary production goes to the benthos. A shift from an ice-associated bloom to a water-column bloom in the central and northern Bering Sea shelf as a result of ice retreat provides the potential for development of the plankton community at the expense of the benthic community[18]. Under climate warming, the benthic community is very likely to be most affected if this carbon is transferred to the deep basin instead of the shelf. Under these circumstances, carbon is disconnected from the food web and can be buried. In contrast, the Barents Sea shelf is much deeper (300 m) and primary production supports a large pelagic community that is unlikely to be affected. Nevertheless, a larger quantity of carbon is likely to be buried in future as deposition shifts from the shelf region to the deeper slope and basin region due to the northward movement of the ice edge.

Projections that the Arctic Ocean will be ice-free in summer (see section 9.2.5.2) imply that production will increase in waters where it was previously limited by ice cover. Based on nutrient availability, Anderson L. et al.[19] estimated that the biological carbon sink would increase by 20 x 1012 g C/yr under ice-free conditions. However, mesocosm studies on the effect of high initial ambient CO2 (750 ?atm) on coccolithophore assemblages have shown an increase in POC production[20]. This would be a negative response to atmospheric CO2 increase.

Alkalinity pump (9.5.3)

Removal of carbonate ions during the formation of calcareous shells and the subsequent sinking of these shells is important in the transfer of inorganic carbon to deeper waters and eventually the sediments. Carbonate shell sinking is also an efficient means of removing organic carbon from the euphotic zone (see Biological Pump above). Together, these processes will provide a negative feedback. However, calcification results in an increase in oceanic pCO2 through the redistribution of carbonate species, which represents a positive feedback. Partial equilibrium with the atmospheric CO2 will result in an increase in pH that may reduce calcification[21].

Terrestrial and coastal sources (9.5.4)

The Arctic Ocean accounts for 20% of the world’s continental shelves and these receive, transport, and store terrestrial organic carbon (primarily from rivers and coastal erosion sources) to an extent significant at the global scale[22]. Olsson and Anderson[23] estimated that 33 to 39 x 1012 g of inorganic carbon are delivered to the Arctic Ocean each year by rivers. Although the amount of total organic carbon is more difficult to estimate because more than 90% is deposited in deltas[24], it may be similar. An increase in precipitation due to climate warming will not necessarily increase carbon burial, however, as the geological composition of the drainage basin and the amount of flow are both controlling factors. For example, the Mackenzie and Yukon are both erosional rivers, while the Siberian rivers are depositional, especially the Ob for which the drainage basin includes marsh lowlands[25]. Thus, increased precipitation is likely to lead to increased DIC delivery in the first case but not the second, and depends on the timing and intensity of the freshwater flow into the sea. Burial will occur on the shelf, and in adjacent ocean basins if transported offshore by sea ice, ocean currents, or turbidity currents.

Regional transport of terrestrial organic carbon to the marine system also results from coastal erosion. For example, the near-shore zone of the Laptev and East Siberian Seas is the most climatically sensitive area in the Arctic and has the highest rates of coastal retreat[26]. Biodegradation of this coastal material is a regional source of high pCO2 in surface waters of the Laptev and East Siberian Seas[27]. Longer ice-free conditions and late-summer storms may accelerate the release of terrestrial carbon frozen during the last glaciation. Pleistocene permafrost soils contain huge ice wedges (up to 60 to 70% by volume) and are enriched by organic carbon (~1 to 20% by weight[28]). The amount of organic carbon stored in permafrost is large (~450 Gt C), similar to the quantity of dissolved carbon stored in the Arctic Ocean[29], and its release to the atmosphere depends on sediment burial rates and competing consumption by biota. The rate of coastal erosion in the Arctic appears to have increased from a few meters per year to tens of meters per year[30]. The highest rates of coastal retreat have been observed at capes; regions important as hunting locations. Bottom erosion is also evident. The bottom depth in the near-shore zone of the Northern Sea Route has increased by ~0.8 m over the past 14 years[31]. Many climate-related factors affect coastal retreat in the Arctic: permafrost ice content, air temperature, wind speed and direction, duration of open water, hydrology, and sea-ice conditions. In addition to the direct effects of climate change, rates of coastal retreat might also increase indirectly due to wave fetch and storm surge activity. Sea-level rise (~15 cm per 100 years[32]) will further accelerate coastal erosion.

Gas hydrates (9.5.1)

The release of CH4 and CO2 trapped in vast gas-hydrate reservoirs in permafrost is very likely to play a key but largely overlooked role in global climate, particularly as CH4 is 60 times more efficient as a GHG (on a molar basis) than CO2. For example, Semiletov[33] estimated that the upper 100 m layer of permafrost contains at least 100,000 Gt of organic carbon in the form of CH4 and CO2. Although CH4 is one of the most important GHGs, there are currently only ~4 Gt of CH4 carbon in the atmosphere. If a small percentage of CH4 from the gas-hydrate reservoir were released to the atmosphere, it could result in an abrupt and significant increase in global temperature through positive feedback effects[34].

The marine Arctic is a particularly important source region for CH4. Following glacial melting and sea-level rise during the Holocene, relatively warm (0°C) Arctic Ocean waters flooded the relatively cold (-12°C) Arctic permafrost domain[35]. As a result, permafrost sediments underlying the arctic shelf regions are still undergoing a dramatic thermal regime change as this heat is conducted downward as a thermal pulse. Subsurface temperatures within the sediment may have risen to the point that both gas hydrate and permafrost may have begun to thaw. In this case CH4 would undergo a phase change, from a stable gas hydrate to a gas, and therefore rise through the sediment. Little is known about the fate of CH4 released in this manner. Depending on the structure and ice matrix of surrounding sediments, CH4 can be either consumed by anaerobic CH4 oxidation or released upward through conduits into the overlying seawater. Evidence of elevated CH4 concentrations in seawater has been observed in the Beaufort Sea[36] and along the North Slope of Alaska[37]. Kvenvolden et al.[38] noted that CH4 concentrations under sea ice in the Beaufort Sea were 3 to 28 times higher in winter than summer, suggesting that CH4 accumulates under the sea ice in winter and is rapidly released into the atmosphere when the sea ice retreats. The timing and release of these under-ice accumulations will change with changes in ice cover.

 

Chapter 9: Marine Systems
9.1. Introduction
9.2. Physical oceanography
    9.2.1. General features
    9.2.2. Sea ice
    9.2.3. Ocean processes of climatic importance
    9.2.4. Variability in hydrographic properties and currents
    9.2.5. Anticipated changes in physical conditions
9.3. Biota
    9.3.1. General description of the Arctic biota community
    9.3.2. Physical factors mediating ecological change
    9.3.3. Past variability – interannual to decadal
    9.3.4. Future change – processes and impacts on biota
9.4. Effects of changes in ultraviolet radiation
9.5. The carbon cycle and climate change
9.6. Key findings
9.7. Gaps in knowledge and research needs

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Glossary

Citation

Committee, I. (2012). Carbon cycle and climate change in the Arctic. Retrieved from http://www.eoearth.org/view/article/150924

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