Over large areas of the surface ocean, the levels
of the nutrients needed for phytoplankton growth
remain high throughout the year. One of the reasons that these nutrients are not exhausted by primary production is lack of micronutrients (such as
iron) for which a source is dust carried from the
continents by winds. Recent iron-enrichment experiments in the Pacific and Southern Oceans have
demonstrated that phytoplankton growth can be
stimulated by small additions of iron (Coale et al.,
1996; Boyd et al., 2000).
1.2.2.5 Carbon dioxide exchanges
The ocean is a major reservoir of carbon dioxide.
In spite of the fact that it already contains 50 times
the mass of CO 2 in the atmosphere, the ocean has
the capacity for storing much more. The surface
ocean exchanges CO 2 with the atmosphere at a
rate proportional to the difference in the partial
pressure of the gas in the two media, with the
ocean able to take up more CO 2 at lower temperatures. The factors limiting the exchange are the
rate at which the gas can be transferred across the
air–sea interface and the rate at which carbon is
exchanged between the surface layer and the deep
ocean. Since ocean currents can transport CO 2
over large distances, the location of storage of carbon may be distant from the site of air–sea
exchange (Wallace, Chapter 6.3).
Photosynthesis by phytoplankton converts CO 2
within the upper ocean to organic compounds.
This lowers the partial pressure of CO 2 in the surface waters, encouraging greater exchange from
the atmosphere. Part of this ‘fixed’ carbon is converted to dissolved inorganic carbon within the
surface layer through respiration and grazing as
well as death and decay. Some of this fixed carbon
leaves the surface layer by particle sinking and by
advection. A small percentage reaches the ocean
floor and is sequestered within ocean sediments
for geological time scales. These issues have been
explored in the Joint Global Ocean Flux Study
(JGOFS) (http://ads.smr.uib.no/jgofs/jgofs.htm).
Annually about 2 Gt of the 6 Gt C released
by burning fossil fuels, is sequestered in the
ocean (Prentice et al., 2001). However, different
approaches reveal different spatial distributions of
this uptake (Wallace, Chapter 6.3). While the rate
at which the ocean stores CO 2 should increase as
the atmospheric concentration rises, changes in
ocean circulation, temperature and/or biological
processes could have significant, but as yet poorly
understood, impacts on the ocean uptake (Sarmiento
et al., 1998; Matear and Hirst, 1999).
1.2.2.6 Effects of sea ice cover
All air–sea exchange processes are greatly modified
in the presence of sea ice, especially when that sea
ice is snow covered. The many roughness elements
on the surface of sea ice transfer more momentum
from the atmosphere to the ice than would occur
between atmosphere and ocean. Roughness elements on the underside of moving sea ice lead to
large transfers of momentum between ice and
ocean. These two factors mean that the net momentum transfer for atmosphere–ice–ocean coupling is
generally greater than for direct atmosphere–ocean
coupling. Sea ice is, however, seldom free to move
and, under a convergent wind stress, much of
the momentum imparted to an ice field is taken up
by internal ice processes, including the formation
of ice ridges and the grounding of ice floes on
shorelines.
Ice is a poor conductor of heat; snow is even
poorer. The rate of heat loss from an ice-covered
ocean is limited by the rate at which heat can be
conducted through the ice/snow layer. In winter,
the presence of ice cover greatly reduces the total
cooling of the water column. However, gaps in the
ice (leads and polynyas) can allow large exchanges
of heat when air–sea temperature differences
are large. Offshore winds can rapidly remove sea
ice between the shore-fast ice and the offshore
pack. Intense ocean cooling and sea ice formation
occurs in these latent heat polynyas (Kaempf and
Backhaus, 1998; Worby et al., 1998).
Newly formed sea ice has a salinity of 6 to 10,
depending on the rate at which it forms. Salts are
concentrated in small pockets of brine that gradually drain into the water column, increasing its
density. Thus the process of sea ice formation
releases salt through brine rejection and creates
deep winter mixed layers through salinity-driven
convection. When the ice melts in the spring, a
low-salinity surface layer is formed, restratifying
the water column (Tang, 1991).
Ice is a good reflector of solar radiation, particularly when the sun’s elevation is low and the ice
is snow covered. During spring at high latitudes,
the long days have the potential to deliver heat to
the surface but the greater part is reflected by ice
and snow cover. When solar elevations are high
SECTION 1 THE OCEAN AND CLIMATE
16
of the nutrients needed for phytoplankton growth
remain high throughout the year. One of the reasons that these nutrients are not exhausted by primary production is lack of micronutrients (such as
iron) for which a source is dust carried from the
continents by winds. Recent iron-enrichment experiments in the Pacific and Southern Oceans have
demonstrated that phytoplankton growth can be
stimulated by small additions of iron (Coale et al.,
1996; Boyd et al., 2000).
1.2.2.5 Carbon dioxide exchanges
The ocean is a major reservoir of carbon dioxide.
In spite of the fact that it already contains 50 times
the mass of CO 2 in the atmosphere, the ocean has
the capacity for storing much more. The surface
ocean exchanges CO 2 with the atmosphere at a
rate proportional to the difference in the partial
pressure of the gas in the two media, with the
ocean able to take up more CO 2 at lower temperatures. The factors limiting the exchange are the
rate at which the gas can be transferred across the
air–sea interface and the rate at which carbon is
exchanged between the surface layer and the deep
ocean. Since ocean currents can transport CO 2
over large distances, the location of storage of carbon may be distant from the site of air–sea
exchange (Wallace, Chapter 6.3).
Photosynthesis by phytoplankton converts CO 2
within the upper ocean to organic compounds.
This lowers the partial pressure of CO 2 in the surface waters, encouraging greater exchange from
the atmosphere. Part of this ‘fixed’ carbon is converted to dissolved inorganic carbon within the
surface layer through respiration and grazing as
well as death and decay. Some of this fixed carbon
leaves the surface layer by particle sinking and by
advection. A small percentage reaches the ocean
floor and is sequestered within ocean sediments
for geological time scales. These issues have been
explored in the Joint Global Ocean Flux Study
(JGOFS) (http://ads.smr.uib.no/jgofs/jgofs.htm).
Annually about 2 Gt of the 6 Gt C released
by burning fossil fuels, is sequestered in the
ocean (Prentice et al., 2001). However, different
approaches reveal different spatial distributions of
this uptake (Wallace, Chapter 6.3). While the rate
at which the ocean stores CO 2 should increase as
the atmospheric concentration rises, changes in
ocean circulation, temperature and/or biological
processes could have significant, but as yet poorly
understood, impacts on the ocean uptake (Sarmiento
et al., 1998; Matear and Hirst, 1999).
1.2.2.6 Effects of sea ice cover
All air–sea exchange processes are greatly modified
in the presence of sea ice, especially when that sea
ice is snow covered. The many roughness elements
on the surface of sea ice transfer more momentum
from the atmosphere to the ice than would occur
between atmosphere and ocean. Roughness elements on the underside of moving sea ice lead to
large transfers of momentum between ice and
ocean. These two factors mean that the net momentum transfer for atmosphere–ice–ocean coupling is
generally greater than for direct atmosphere–ocean
coupling. Sea ice is, however, seldom free to move
and, under a convergent wind stress, much of
the momentum imparted to an ice field is taken up
by internal ice processes, including the formation
of ice ridges and the grounding of ice floes on
shorelines.
Ice is a poor conductor of heat; snow is even
poorer. The rate of heat loss from an ice-covered
ocean is limited by the rate at which heat can be
conducted through the ice/snow layer. In winter,
the presence of ice cover greatly reduces the total
cooling of the water column. However, gaps in the
ice (leads and polynyas) can allow large exchanges
of heat when air–sea temperature differences
are large. Offshore winds can rapidly remove sea
ice between the shore-fast ice and the offshore
pack. Intense ocean cooling and sea ice formation
occurs in these latent heat polynyas (Kaempf and
Backhaus, 1998; Worby et al., 1998).
Newly formed sea ice has a salinity of 6 to 10,
depending on the rate at which it forms. Salts are
concentrated in small pockets of brine that gradually drain into the water column, increasing its
density. Thus the process of sea ice formation
releases salt through brine rejection and creates
deep winter mixed layers through salinity-driven
convection. When the ice melts in the spring, a
low-salinity surface layer is formed, restratifying
the water column (Tang, 1991).
Ice is a good reflector of solar radiation, particularly when the sun’s elevation is low and the ice
is snow covered. During spring at high latitudes,
the long days have the potential to deliver heat to
the surface but the greater part is reflected by ice
and snow cover. When solar elevations are high
SECTION 1 THE OCEAN AND CLIMATE
16