188.8.131.52 Future changes in ocean CO2 uptake
This section lists processes that may be important for the future uptake of
anthropogenic CO2. These changes can represent changes in anthropogenic
CO2 uptake itself (mainly physical and chemical processes), or changes
in the natural biologically-linked cycling of carbon between the atmosphere
Physical and chemical processes
Buffering changes. The capacity of surface waters to take up anthropogenic
CO2 is decreasing as CO2 levels increase (see Box
3.3). The magnitude of this effect is substantial. This decrease in uptake
capacity of the ocean makes atmospheric CO2 more sensitive to anthropogenic
emissions and other changes in the natural cycling of carbon.
Emissions rate. Even assuming no other changes to the carbon cycle, the
proportion of emitted CO2 that can be taken up by the ocean decreases
as the rate of emission increases. This is due to the finite rate of exposure
of older’, deeper waters to the anthropogenic CO2 contained
in the atmosphere.
Warming. CO2 is less soluble in warmer water, and the equilibrium
pCO2 in seawater increases by about 10 to 20 ppm per °C
temperature increase. Warming of surface water would therefore tend to increase
surface water pCO2, driving CO2 from the surface
ocean to the atmosphere. The expected effect of such warming on atmospheric
CO2 may be smaller, depending on the rate of exchange between ocean
surface waters and the deep ocean at high latitudes (e.g., Bacastow, 1993).
Vertical mixing and stratification. Several coupled atmosphere-ocean
models have shown global warming to be accompanied by an increase in vertical
stratification (see Chapter 7). Such a change would reduce
the rate of mixing between surface and deep waters, and therefore reduce the
effective volume of the ocean that is exposed to high atmospheric CO2.
On its own, this effect would tend to reduce the ocean CO2 uptake.
However, changes in stratification may also drive changes in the natural carbon
cycle. The magnitude and even the sign of changes in the natural cycle are much
more difficult to predict because of the complexity of ocean biological processes
(Sarmiento et al., 1998; Matear and Hirst, 1999).
Qualitative and quantitative changes in carbon uptake arising from changes in
marine ecosystems are more speculative (Denman et al., 1996; Falkowski et al.,
1998; Watson and Liss, 1998), but are likely to have occurred over glacial-interglacial
time-scales (Section 3.3). Falkowski et al. (1998) listed
three major classes of biologically linked factors that can in principal alter
the air-sea partitioning of CO2: (1) changes in surface nutrient
utilisation (e.g., in HNLC areas); (2) changes in total ocean content of major
nutrients; (3) changes in the elemental composition of biogenic material (including
the rain ratio). Our incomplete understanding of present day nutrient controls
on productivity limits our ability to predict future changes in ocean biology
and their effect on CO2 levels. For example, the possible identification
of changes in deep ocean C:N:P ratios (Pahlow and Riebesell, 2000) leaves open
the question of the extent to which ocean biological carbon cycling is in steady
state, or is likely to remain so in the future.
Changes in surface nutrient utilisation. Changes in the utilisation
of surface nutrients in HNLC regions have the potential to alter export production
and carbon storage in the ocean interior. Most attention focuses on the role
of inadvertent or deliberate changes in the external supply of iron to such
regions. The sign of possible future responses of ocean biota due to iron supply
changes is difficult to assess. Future iron supply may increase due to erosion
(enhanced by agriculture and urbanisation) which tends to increase dust export
and aeolian iron deposition (Tegen and Fung, 1995). Conversely, a globally enhanced
hydrological cycle and increased water-use efficiency of terrestrial plants
may tend to reduce future dust export (Harrison et al., 2001). The delivery
of dust to the HNLC regions will be sensitive to regional changes in erosion
and the hydrological cycle, affecting the important regions of dust export,
rather than to global scale changes (Dai et al., 1998).
Surface nutrient supply could be reduced if ocean stratification reduces the
supply of major nutrients carried to the surface waters from the deep ocean
(Sarmiento et al., 1998). The impact of stratification on marine productivity
depends on the limiting factor. In regions limited by deep ocean nutrients,
stratification would reduce marine productivity and the strength of the export
of carbon by biological processes. Conversely, stratification also increases
the light exposure of marine organisms, which would increase productivity in
regions where light is limiting.
Changes in total ocean content of major nutrients. Changes in the delivery
of the major biologically limiting nutrients (N, P, Fe, Si) from riverine, atmospheric
or sedimentary sources, or changes in removal rates (e.g., denitrification),
could alter oceanic nutrient inventories and hence export production and ocean
carbon storage. On the global scale, the upward fluxes of major nutrients are
slightly depleted in N relative to P with respect to the nutrient requirements
of phytoplankton (Fanning, 1992). This relative supply of N versus P may be
sensitive to climate and circulation related changes in the rate of fixed-nitrogen
removal by denitrification (Ganeshram et al., 1995) or via changes in the rate
of nitrogen fixation. Changes in river flow and composition are also affecting
the supply of nutrients (Frankignoulle et al., 1998). The hypothesised link
between nitrogen fixation in certain ocean regions and the external iron supply
(Falkowski, 1997; Wu et al., 2000) could play a role in future nutrient and
carbon budgets. Nitrogen fixation rates may also be affected by changes in stratification
and mixing. For example, Karl et al. (1997) have identified interannual variability
in nitrogen fixation rates in the sub-tropical Pacific which are apparently
linked to ENSO variability in upper ocean dynamics.
Changes in the elemental composition of biogenic material. The structure
and biogeochemistry of marine ecosystems can be affected by numerous climate-related
factors including temperature, cloudiness, nutrient availability, mixed-layer
physics and sea-ice extent. In turn the structure of marine ecosystems, and
particularly the species composition of phytoplankton, exert a control on the
partitioning of carbon between the ocean and the atmosphere. For example, a
change in distribution of calcareous versus siliceous planktonic organisms could
affect CO2 uptake in the future, as it may have done in the past
(Archer and Maier-Reimer, 1994). Precipitation of CaCO3 by marine
organisms (calcification) removes dissolved CO32-, thus
decreasing surface water alkalinity and reducing the capacity of sea water to
dissolve atmospheric CO2 (see Box 3.3).
Recent experimental evidence suggests that as a direct result of increasing
atmospheric and surface water pCO2 levels, oceanic calcification
will decrease significantly over the next 100 years. Model-based calculations
suggest that decreases in coral reef calcification rates of the order 17 to
35% relative to pre-industrial rates are possible (Kleypas et al., 1999). Experimental
studies with corals have confirmed such effects (Langdon et al., 2000). Field
and laboratory studies have shown that planktonic calcification is also highly
sensitive to pCO2 levels. The calcification rate of coccolithophorids
decreases by 16 to 83% at pCO2 levels of 750 ppm (Riebesell
et al., 2000). Such an effect would tend to favour CO2 storage in
the upper ocean and act as a negative feedback on atmospheric growth rates of
CO2. However, long-term predictions of such biological responses
are hampered by a lack of understanding concerning physiological acclimation
and genetic adaptations of species to increasing pCO2.
Box 3.4: Causes of glacial/inter-glacial changes
in atmospheric CO2.
One family of hypotheses to explain glacial/inter-glacial variations
of atmospheric CO2 relies on physical mechanisms that could
change the dissolution and outgassing of CO2 in the ocean.
The solubility of CO2 is increased at low temperature, but
reduced at high salinity. These effects nearly cancel out over the glacial/inter-glacial
cycle, so simple solubility changes are not the answer. Stephens and Keeling
(2000) have proposed that extended winter sea ice prevented outgassing
of upwelled, CO2-rich water around the Antarctic continent
during glacial times. A melt-water “cap” may have further restricted
outgassing of CO2 during summer (François et al., 1997).
These mechanisms could explain the parallel increases of Antarctic temperature
and CO2 during deglaciation. However, they require less vertical
mixing to occur at low latitudes than is normally assumed. The relative
importance of high and low latitudes for the transport of CO2
by physical processes is not well known, and may be poorly represented
in most ocean carbon models (Toggweiler, 1999; Broecker et al., 1999).
Several authors have hypothesised increased utilisation of surface nutrients
by marine ecosystems in high latitudes, leading to stronger vertical gradients
of DIC and thus reduced atmospheric CO2 during glacial times
(Sarmiento and Toggweiler, 1984; Siegenthaler and Wenk, 1984; Knox and
McElroy, 1984). Other hypotheses call for an increased external supply
of nutrients to the ocean (McElroy, 1983; Martin et al., 1990; Broecker
and Henderson, 1998). The supply of iron-rich dust to the Southern Ocean
is increased during glacial periods, due to expanded deserts in the Patagonian
source region (Andersen et al., 1998; Mahowald et al., 1999; Petit et
al., 1999); dust-borne iron concentration in Antarctic ice is also increased
(Edwards et al., 1998). Fertilisation of marine productivity by iron from
this source could have influenced atmospheric CO2. Most of
these mechanisms, however, can only account for about 30 ppm, or less,
of the change (Lefèvre and Watson, 1999; Archer and Johnson, 2000).
Palaeo-nutrient proxies have also been used to argue against large changes
in total high latitude productivity (Boyle, 1988; Rickaby and Elderfield,
1999; Elderfield and Rickaby, 2000), even if the region of high productivity
in the Southern Ocean may have been shifted to the north (Kumar et al.,
1995; François et al., 1997). Increased productivity over larger
regions might have been caused by decreased denitrification (Altabet et
al., 1995; Ganeshram et al., 1995) or iron stimulated N2 fixation
(Broecker and Henderson, 1998) leading to an increase in the total ocean
content of reactive nitrogen.
Another family of hypotheses invokes ocean alkalinity changes by a variety
of mechanisms (Opdyke and Walker, 1992; Archer and Maier-Reimer, 1994;
Kleypas, 1997), including increased silica supply through dust, promoting
export production by siliceous rather than calcareous phytoplankton (Harrison,
2000). Although there is geochemical evidence for higher ocean pH during
glacial times (Sanyal et al., 1995), a large increase in alkalinity would
result in a much deeper lysocline, implying an increase in CaCO3
preservation that is not observed in deep-sea sediments (Catubig et al.,
1998; Sigman et al., 1998; Archer et al., 2000).
Given the complex timing of changes between climate changes and atmospheric
CO2 on glacial-interglacial time-scales, it is plausible that
more than one mechanism has been in operation; and indeed most or all
of the hypotheses encounter difficulties if called upon individually to
explain the full magnitude of the change.