Past changes in ocean carbonate chemistry - Ocean Acidification - page 26

important for understanding changes in surface-
ocean chemistry and atmospheric CO 2 on glacial-
interglacial timescales. While surface-ocean changes
during the glacial-interglacial cycles were ~80 μatm
in p CO 2 and ~0.2 units in pH, deep-ocean carbonate
chemistry changes were probably much smaller
(see Section 2.4.2 and Zeebe and Marchitto 2010).

inl ux of Ca 2+ and CO 3 2- over burial of CaCO 3 . Then,
the concentrations of Ca 2+ and CO 3 2- in seawater
increase which leads to an increase in the CaCO 3
saturation state. This in turn leads to a deepening of
the saturation horizon and to an increased burial of
CaCO 3 until the burial again balances the inl ux.
The new balance is restored at higher CO 3 2- than
before.

2.3.4 Tectonic (>100 000 yr) timescale

2.3.6 Biological pump

A large amount of carbon is locked up in the earth's
crust as carbonate carbon (~70 × 10 6 Pg C) and as
elemental carbon in shales and coals (~20 × 10 6
Pg C). On timescales >100 000 yr, this reservoir is
active, and imbalances in the l uxes to and from this
pool can lead to large changes in C T , A T , and atmos-
pheric CO 2 (Fig. 2.1B). The balance between CO 2
consumption by subduction of marine sediments,
net organic carbon burial, weathering, subsequent
carbonate burial, and volcanic degassing of CO 2 is
the dominant process controlling carbon l uxes on
this timescale (e.g. Walker et al. 1981 ; Berner et al.
1983 ; Caldeira 1992 ; Zeebe and Caldeira 2008 ).
Figure 2.1 illustrates the fundamental difference
between short-term carbon cycling on, for example,
a 10 to 100 yr timescale (Fig. 2.1A) and long-term
carbon cycling (Fig. 2.1B). The two distinct cycles
involve vastly different reservoir sizes and different
sets of controls on atmospheric CO 2 and ocean
chemistry. It follows that carbon cycling and ocean
chemistry conditions during long-term steady states
(e.g. over millions of years) cannot simply be com-
pared with rapid, transient events (e.g. over the
next few centuries). Section 2.4.5 provides more
details.

In nature, C T and A T are manipulated both by biotic
and by abiotic factors. For instance, the removal of
dissolved inorganic carbon from the ocean by
phytoplankton at the ocean surface and the creation
of carbon-based cellular organic components
(carbohydrates, fats, proteins, etc.) drives down the
CO 2 concentration. If the cells were subsequently
degraded in the surface layer and compounds bro-
ken down back into their basic components, the C T
concentration would be restored and atmospheric
p CO 2 unaffected. The net effect is zero if aerobic
respiration balances photosynthetic production,
because respiration and photosynthesis run the
same equation in opposite directions (see Chapter
1). It does not quite happen like this in the real
ocean, as on average about 10% of primary produc-
tion (and C T ) removal escapes the surface layer and
settles gravitationally into the ocean interior before
being broken down ('remineralized'). Ultimately,
ocean circulation works to bringing the excess C T
back to the surface, but a gradient is created with
higher C T at depth. Hence, the faster the rate of
export of particulate organic carbon (POC) from the
surface ocean, or the deeper it can sink without
being degraded, the stronger the C T gradient, the
lower the surface-ocean CO 2 concentrations, and
hence the lower the atmospheric p CO 2 . This process
is known loosely as the 'biological pump'.
Past changes in the strength of the biological
pump have thus modulated the C T concentration at
the surface, and by inference the acidity (pH). For
instance, it has been hypothesized that during the
last glacial period, the strength of the biological
pump was greater, meaning lower atmospheric
p CO 2 and higher pH. Reconstructions of changes in
ocean-surface pH based on the boron isotopic com-
position of marine carbonates (boron speciation in

2.3.5 Calcite compensation

Calcite compensation maintains the balance
between CaCO 3 weathering l uxes into the ocean
and CaCO 3 burial l uxes in marine sediments on a
timescale of 5000 to 10 000 years (e.g. Broecker and
Peng 1989 ; Zeebe and Westbroek 2003 ). At steady
state, the riverine l ux of Ca 2+ and CO 3 2- ions from
weathering must be balanced by burial of CaCO 3 in
the sea, otherwise [Ca 2+ ] and [CO 3 2- ] would rise or
fall. The feedback that maintains this balance works
as follows. Assume there is an excess weathering