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12.2 Ocean Acidification and Its Effect on Photosynthetic
Organisms
12.2.1 Ocean Acidification
The exchange of CO 2 between the oceans and the atmosphere relies on seawater
mixing intensity and surface ocean carbonate chemistry. The dissolved CO 2 in
seawater remains in the upper layer for 6 years on average. The mixture between
epipelagic and mesopelagic seawater (1000–4000 m) is relatively slow and needs
hundreds of years (The Royal Society 2005). The oceans have absorbed more than
1/3 of the anthropogenically released CO 2 since the Industrial Revolution, which
significantly mitigates global warming (Sabine et al. 2004). However, the atmospheric CO 2 concentration is still rising, with continuous oceanic absorption of CO 2 ,
leading to a lowered alkalinity of the mesopelagic seawater. This process of increasing acidity of seawater caused by the rising atmospheric CO 2 concentration is
termed “ocean acidification (OA).” More than half of the CO 2 absorbed from the
atmosphere by the oceans remains at a depth of 0 to 400 m (Sabine et al. 2004); the
longer CO 2 remains in the upper layer, the faster OA proceeds.
CO 2 is an acidic gas, and its dissolution into seawater leads to acidification:
CO H O H CO
2
2
2
3
+
=
(12.1)
H CO H HCO
2
3
3
=
+
+
–
(12.2)
HCO
H CO
3
3
2
−
+
=
+
–
(12.3)
(The equilibrium constants of the above three reactions are dependent on temperature and salinity in surface oceans). As indicated by these reactions, when CO 2
dissolves into seawater, it forms carbonic acid. Then, carbonic acid dissociates and
forms HCO 3
− and H
+
. As the concentration of H
+
increases, reaction (12.3) is shifted
to the right, causing the decrease of CO 3
2− concentration.
The seawater carbonate system of the upper oceans provides an inorganic carbon
source for primary production. With the progressive pH decrease associated with
OA, the concentration of inorganic carbon in the carbonate system and the concentration ratios of different inorganic carbons (CO 2 , HCO 3
− , CO 3
2− ) change, thus
affecting the saturation Ω of CaCO 3 in seawater (Ω = Ca
2+
× CO 3
2− /K c , where K c is
the product of Ca
2+
× CO 3
2− when the CaCO 3 solution is saturated, which is associated with the crystal type of CaCO 3 such as calcite and aragonite. Since the oceanic
Ca
2+
concentration is relatively constant (approximately 10 mM), the CaCO 3 saturation Ω mainly depends on the concentration of CO 3
2− . In general, HCO 3
− in the
seawater accounts for more than 90% of the dissolved inorganic carbon (DIC),
CO 3
2− for about 9%, and CO 2 for less than 1% (these percentages change at different latitudes or regions). Increasing atmospheric CO 2 results in increased concentrations of dissolved CO 2 , HCO 3
− , and H
+
, the decreased concentration of CO 3
2− , and
12 Effects of Ocean Acidification and UV Radiation on Marine Photosynthetic…
12.2 Ocean Acidification and Its Effect on Photosynthetic
Organisms
12.2.1 Ocean Acidification
The exchange of CO 2 between the oceans and the atmosphere relies on seawater
mixing intensity and surface ocean carbonate chemistry. The dissolved CO 2 in
seawater remains in the upper layer for 6 years on average. The mixture between
epipelagic and mesopelagic seawater (1000–4000 m) is relatively slow and needs
hundreds of years (The Royal Society 2005). The oceans have absorbed more than
1/3 of the anthropogenically released CO 2 since the Industrial Revolution, which
significantly mitigates global warming (Sabine et al. 2004). However, the atmospheric CO 2 concentration is still rising, with continuous oceanic absorption of CO 2 ,
leading to a lowered alkalinity of the mesopelagic seawater. This process of increasing acidity of seawater caused by the rising atmospheric CO 2 concentration is
termed “ocean acidification (OA).” More than half of the CO 2 absorbed from the
atmosphere by the oceans remains at a depth of 0 to 400 m (Sabine et al. 2004); the
longer CO 2 remains in the upper layer, the faster OA proceeds.
CO 2 is an acidic gas, and its dissolution into seawater leads to acidification:
CO H O H CO
2
2
2
3
+
=
(12.1)
H CO H HCO
2
3
3
=
+
+
–
(12.2)
HCO
H CO
3
3
2
−
+
=
+
–
(12.3)
(The equilibrium constants of the above three reactions are dependent on temperature and salinity in surface oceans). As indicated by these reactions, when CO 2
dissolves into seawater, it forms carbonic acid. Then, carbonic acid dissociates and
forms HCO 3
− and H
+
. As the concentration of H
+
increases, reaction (12.3) is shifted
to the right, causing the decrease of CO 3
2− concentration.
The seawater carbonate system of the upper oceans provides an inorganic carbon
source for primary production. With the progressive pH decrease associated with
OA, the concentration of inorganic carbon in the carbonate system and the concentration ratios of different inorganic carbons (CO 2 , HCO 3
− , CO 3
2− ) change, thus
affecting the saturation Ω of CaCO 3 in seawater (Ω = Ca
2+
× CO 3
2− /K c , where K c is
the product of Ca
2+
× CO 3
2− when the CaCO 3 solution is saturated, which is associated with the crystal type of CaCO 3 such as calcite and aragonite. Since the oceanic
Ca
2+
concentration is relatively constant (approximately 10 mM), the CaCO 3 saturation Ω mainly depends on the concentration of CO 3
2− . In general, HCO 3
− in the
seawater accounts for more than 90% of the dissolved inorganic carbon (DIC),
CO 3
2− for about 9%, and CO 2 for less than 1% (these percentages change at different latitudes or regions). Increasing atmospheric CO 2 results in increased concentrations of dissolved CO 2 , HCO 3
− , and H
+
, the decreased concentration of CO 3
2− , and
12 Effects of Ocean Acidification and UV Radiation on Marine Photosynthetic…