724
K. M. G. Mostofa et al.
contents of organic matter (DOM and POM) in coastal waters are responsible for
the higher observed contents of Chl a compared to the open ocean (Clark et al.
2004). It is generally known that DOM and POM (e.g. phytoplankton) can release
NO 3
− and PO 4
3− , by either photoinduced or microbial assimilation/respiration in
waters (see chapters “Dissolved Organic Matter in Natural Waters, Photoinduced
and Microbial Degradation of Dissolved Organic Matter in Natural Waters” and
“Photosynthesis in Nature: A New Look”).
(ii) Strong wind and wave mixing along with the solar (UV and PAR) radiation
may degrade Chl a, DNA or biomolecules bound to PSI and PSII of microorganisms. The effect would be more marked in the open ocean compared to coastal
waters. This issue would be supported by the observation that UV-B radiation
(280–315 nm) can inhibit photosynthetic carbon fixation by tropical phytoplankton
assemblages in coastal to pelagic surface seawaters (Li et al. 2011). The inhibition
of photosynthesis by UV-A (315–400 nm) increases from coastal to offshore waters
(Li et al. 2011). It has also been shown that UV-B inhibits photosynthesis by up to
27 % and UV-A by up to 29 % (Li et al. 2011). In East China Sea, lower concentration of Chl a (0.06–0.07 μg L −1 : Kuroshio sites) has been detected in the open
ocean, with high water temperature (23.9–24.0 °C) and low NO 3
− (<0.1 μM), than
in coastal seawater (0.43–2.44 μg L −1 ) (Hung et al. 2000). The latter had low temperature (16.3–18.9 °C) and high NO 3
− (<0.4–6.0 μM) (Hung et al. 2000).
Similarly and interestingly, Chl a concentrations are largely variable (0.06–
1,000 μg L −1 ) and substantially high (occasionally >1,000 μg L −1 ) in ice- covered
Antarctic and Arctic Oceans (Table 1) (Palmisano et al. 1985; Garrison et al. 1986;
Wheeler et al. 1996; Mock and Gradinger 1999; Apollonio 1980; Guildford and
Hecky 2000; Norrbin et al. 2009; Sakshaug and Holm-Hansen 1986; Spies 1987;
Verlencar et al. 1990; Varela et al. 2002; Cottrell and Kirchman 2009; Hewes et
al. 2009). The highest Chl a concentrations, reaching values higher than 1,000 μg
L −1 , have been detected in bottom-ice communities of Antarctica Ocean.
Otherwise, Chl a is largely variable: it reached <297 μg L −1 in the ice undersurface; <5.2 in the water column of central Arctic Ocean; <86 μg L −1 in Barents and
Greenland Sea ice (Arctic Ocean); <25 μg L −1 in Gerlache and south Bransfield
Straits (Antarctic Peninsula); <8.2 μg L −1 in Dumbell Bay (Arctic Ocean);
<4.03 μg L −1 in ocean seawater (Antarctic Ocean); <4.0 μg L −1 in South Shetland
Islands (Antarctica), 0.10–2.27 in other ice seawater; and finally 111 ± 30 μg L −1
in incubation experimental studies using Antarctic ice seawater (Table 1).
It has been shown that Chl a varies significantly, from 0.1 to 297 μg L −1 in ice
undersurface and from 0.1 to 5.2 μg L −1 in the water column (Wheeler et al. 1996).
The values of Chl a can increase in the range of the potential phytoplankton standing stock (25–50 μg L −1 ) in Antarctic marine waters, Southern Ocean (Sakshaug
and Holm-Hansen 1986; Spies 1987). Similarly, Chl a contents in bottom ice communities reach 300–400 mg m −2 (Steemann-Nielsen 1962; Palmisano and Sullivan
1983). Such numerous algal communities are presumably the consequence of several phenomena: (i) Algal growth may be prolonged due to low temperature and low
solar irradiance, which are unable to form O 2 • − and subsequently H 2 O 2 or HO
• . This
phenomenon can protect algal cells from death, allowing high primary production
K. M. G. Mostofa et al.
contents of organic matter (DOM and POM) in coastal waters are responsible for
the higher observed contents of Chl a compared to the open ocean (Clark et al.
2004). It is generally known that DOM and POM (e.g. phytoplankton) can release
NO 3
− and PO 4
3− , by either photoinduced or microbial assimilation/respiration in
waters (see chapters “Dissolved Organic Matter in Natural Waters, Photoinduced
and Microbial Degradation of Dissolved Organic Matter in Natural Waters” and
“Photosynthesis in Nature: A New Look”).
(ii) Strong wind and wave mixing along with the solar (UV and PAR) radiation
may degrade Chl a, DNA or biomolecules bound to PSI and PSII of microorganisms. The effect would be more marked in the open ocean compared to coastal
waters. This issue would be supported by the observation that UV-B radiation
(280–315 nm) can inhibit photosynthetic carbon fixation by tropical phytoplankton
assemblages in coastal to pelagic surface seawaters (Li et al. 2011). The inhibition
of photosynthesis by UV-A (315–400 nm) increases from coastal to offshore waters
(Li et al. 2011). It has also been shown that UV-B inhibits photosynthesis by up to
27 % and UV-A by up to 29 % (Li et al. 2011). In East China Sea, lower concentration of Chl a (0.06–0.07 μg L −1 : Kuroshio sites) has been detected in the open
ocean, with high water temperature (23.9–24.0 °C) and low NO 3
− (<0.1 μM), than
in coastal seawater (0.43–2.44 μg L −1 ) (Hung et al. 2000). The latter had low temperature (16.3–18.9 °C) and high NO 3
− (<0.4–6.0 μM) (Hung et al. 2000).
Similarly and interestingly, Chl a concentrations are largely variable (0.06–
1,000 μg L −1 ) and substantially high (occasionally >1,000 μg L −1 ) in ice- covered
Antarctic and Arctic Oceans (Table 1) (Palmisano et al. 1985; Garrison et al. 1986;
Wheeler et al. 1996; Mock and Gradinger 1999; Apollonio 1980; Guildford and
Hecky 2000; Norrbin et al. 2009; Sakshaug and Holm-Hansen 1986; Spies 1987;
Verlencar et al. 1990; Varela et al. 2002; Cottrell and Kirchman 2009; Hewes et
al. 2009). The highest Chl a concentrations, reaching values higher than 1,000 μg
L −1 , have been detected in bottom-ice communities of Antarctica Ocean.
Otherwise, Chl a is largely variable: it reached <297 μg L −1 in the ice undersurface; <5.2 in the water column of central Arctic Ocean; <86 μg L −1 in Barents and
Greenland Sea ice (Arctic Ocean); <25 μg L −1 in Gerlache and south Bransfield
Straits (Antarctic Peninsula); <8.2 μg L −1 in Dumbell Bay (Arctic Ocean);
<4.03 μg L −1 in ocean seawater (Antarctic Ocean); <4.0 μg L −1 in South Shetland
Islands (Antarctica), 0.10–2.27 in other ice seawater; and finally 111 ± 30 μg L −1
in incubation experimental studies using Antarctic ice seawater (Table 1).
It has been shown that Chl a varies significantly, from 0.1 to 297 μg L −1 in ice
undersurface and from 0.1 to 5.2 μg L −1 in the water column (Wheeler et al. 1996).
The values of Chl a can increase in the range of the potential phytoplankton standing stock (25–50 μg L −1 ) in Antarctic marine waters, Southern Ocean (Sakshaug
and Holm-Hansen 1986; Spies 1987). Similarly, Chl a contents in bottom ice communities reach 300–400 mg m −2 (Steemann-Nielsen 1962; Palmisano and Sullivan
1983). Such numerous algal communities are presumably the consequence of several phenomena: (i) Algal growth may be prolonged due to low temperature and low
solar irradiance, which are unable to form O 2 • − and subsequently H 2 O 2 or HO
• . This
phenomenon can protect algal cells from death, allowing high primary production