3.2.3.2 Uptake of anthropogenic CO2
Figure 3.2: Variations in atmospheric CO2 concentration on
different time-scales. (a) Direct measurements of atmospheric CO2
concentration (Keeling and Whorf, 2000), and O2 from 1990 onwards
(Battle et al., 2000). O2 concentration is expressed as the change
from an arbitrary standard. (b) CO2 concentration in Antarctic
ice cores for the past millenium (Siegenthaler et al., 1988; Neftel et al.,
1994; Barnola et al., 1995; Etheridge et al., 1996). Recent atmospheric
measurements at Mauna Loa (Keeling and Whorf, 2000) are shown for comparison.
(c) CO2 concentration in the Taylor Dome Antarctic ice core (Indermühle
et al., 1999). (d) CO2 concentration in the Vostok Antarctic
ice core (Petit et al., 1999; Fischer et al., 1999). (e) Geochemically inferred
CO2 concentrations, from Pagani et al. (1999a) and Pearson and
Palmer (2000). (f) Geochemically inferred CO2 concentrations:
coloured bars represent different published studies cited by Berner (1997).
The data from Pearson and Palmer (2000) are shown by a black line. (BP =
before present.) |
Despite the importance of biological processes for the
ocean’s natural carbon cycle, current thinking maintains that the oceanic
uptake of anthropogenic CO2 is primarily a physically and chemically
controlled process superimposed on a biologically driven carbon cycle that is
close to steady state. This differs from the situation on land because of the
different factors which control marine and terrestrial primary productivity.
On land, experiments have repeatedly shown that current CO2 concentrations
are limiting to plant growth (Section 3.2.2.4). In the
ocean, experimental evidence is against control of productivity by CO2
concentrations, except for certain species at lower than contemporary CO2
concentrations (Riebesell et al., 1993; Falkowski, 1994). Further, deep ocean
concentrations of major nutrients and DIC are tightly correlated, with the existing
ratios closely (but not exactly, see Section 3.2.3.3)
matching the nutritional requirements of marine organisms (the “Redfield
ratios”: Redfield et al., 1963). This implies that as long as nutrients
that are mixed into the ocean surface layer are largely removed by organic carbon
production and export, then there is little potential to drive a net air-sea
carbon transfer simply through alteration of the global rate of production.
Terrestrial ecosystems show greater variability in this respect because land
plants have multiple ways to acquire nutrients, and have greater plasticity
in their chemical composition (Melillo and Gosz, 1983). There are, however,
extensive regions of the ocean surface where major nutrients are not fully depleted,
and changes in these regions may play a significant role in altering atmosphere-ocean
carbon partitioning (see Section 3.2.3.3).
The increase of atmospheric pCO2 over pre-industrial levels
has tended to increase uptake into natural CO2 sink regions and decreased
release from natural outgassing regions. Contemporary net air-sea fluxes comprise
spatially-varying mixtures of natural and anthropogenic CO2 flux
components and cannot be equated with anthropogenic CO2 uptake, except
on a global scale. Uptake of anthropogenic CO2 is strongest in regions
where “old” waters, which have spent many years in the ocean interior
since their last contact with the atmosphere, are re-exposed at the sea surface
to a contemporary atmosphere which now contains anthropogenic CO2
(e.g., Sarmiento et al., 1992; Doney, 1999). In an upwelling region, for example,
the natural component of the air-sea flux may be to outgas CO2 to
the atmosphere. The higher atmospheric pCO2 of the contemporary
atmosphere acts to reduce this outgassing relative to the natural state, implying
that more carbon remains in the ocean. This represents uptake of anthropogenic
CO2 by a region which is a source of CO2 to the atmosphere.
The additional carbon in the ocean resulting from such uptake is then transported
by the surface ocean circulation, and eventually stored as surface waters sink,
or are mixed, into the deep ocean interior. Whereas upwelling into the surface
layer is quantitatively balanced on a global scale by sinking, the locations
where deep waters rise and sink can be separated by large horizontal distances.
Air-sea gas transfer allows older waters to approach a new steady state with
higher atmospheric CO2 levels after about a year at the sea surface.
This is fast relative to the rate of ocean mixing, implying that anthropogenic
CO2 uptake is limited by the rate at which “older” waters
are mixed towards the air-sea interface. The rate of exposure of older, deeper
waters is therefore a critical factor limiting the uptake of anthropogenic CO2.
In principle, there is sufficient uptake capacity (see Box 3.3) in the ocean
to incorporate 70 to 80% of anthropogenic CO2 emissions to the atmosphere,
even when total emissions of up to 4,500 PgC are considered (Archer et al.,
1997). The finite rate of ocean mixing, however, means that it takes several
hundred years to access this capacity (Maier-Reimer and Hasselmann, 1987; Enting
et al., 1994; Archer et al., 1997). Chemical neutralisation of added CO2
through reaction with CaCO3 contained in deep ocean sediments could
potentially absorb a further 9 to 15% of the total emitted amount, reducing
the airborne fraction of cumulative emissions by about a factor of 2; however
the response time of deep ocean sediments is in the order of 5,000 years (Archer
et al., 1997).
Using time-series and global survey data, the increasing oceanic carbon content
has been directly observed, although the signal is small compared to natural
variability and requires extremely accurate measurements (Sabine et al., 1997).
A long-term increase of surface water CO2 levels tracking the mean
atmospheric CO2 increase has been observed in the ocean’s subtropical
gyres (Bates et al., 1996; Winn et al., 1998) and the equatorial Pacific (Feely
et al., 1999b). However, very few such time-series exist and the response of
other important oceanic regions to the atmospheric pCO2 increase
cannot yet be assessed. Inter-decadal increases in DIC concentrations at depth
have been resolved from direct measurements (Wallace, 1995; Peng et al., 1998;
Ono et al., 1998; Sabine et al., 1999). The total amounts of anthropogenic CO2
accumulated in the ocean since the pre-industrial era can also be estimated
from measurements using recent refinements (Gruber et al., 1996) of long-standing
methods for separating the natural and anthropogenic components of oceanic DIC.
A comparison of such analyses with ocean model results is presented in Section
3.6.3.
