enhance their abilities to remain suspended in the
upper water column while alive. These same characteristics largely dictate their resistance to dissolution
after death. Taxa living in warm, tropical surface
waters, where density is generally low, tend to be
open-structured with thin shells and porous walls.
Taxa that live deeper in cooler, denser subsurface
waters, or in colder surface waters at high latitudes,
tend to be more heavily calcified with thicker shells
and small or closed up pores. On the seafloor, the thinshelled, more fragile species tend to dissolve more
readily than the robust taxa. In effect, this means that
individual species each have their own ‘lysocline’,
which can be offset shallower or deeper from the
foraminiferal lysocline determined from the total assemblage. There are additional consequences of this
selective preservation of taxa that must be considered
in paleo-oceanographic or paleoclimatic studies. For
example, the selective preservation of more heavily
calcified taxa tends to impart a generally ‘cooler’ appearance to the overall microfossil population and
can bias attempts to derive paleotemperature information from seafloor assemblages, as well as other
population properties such as diversity.
For carbonate particles produced in the upper
ocean, settling rates play an important role in their
distribution and preservation. Smaller planktonic
foraminifers settle at about 150–250 m d
À1 , while
larger (4250 mm) foraminifers may settle as much as
2000 m d
À1
. These rates are rapid enough that little
dissolution is thought to occur in the water column.
Solitary coccoliths, on the other hand, sink at rates of
0.3 to B10 m d
À1
, slow enough that dissolution
within the water column should theoretically prevent
their ever reaching the ocean bottom. However,
sediment trap studies have shown that transport by
fecal pellets is the dominant process by which small
phytoplankton skeletons are transferred to the seafloor. Protection offered by the organic fecal pellet
covering may also protect the coccoliths after deposition and account for the fact that the coccolith
lysocline is generally observed to lie somewhat deeper than the foraminiferal lysocline.
While the seafloor depths of the lysocline and
CCD can be readily identified from sedimentary
criteria, this information is of limited use without
realistic knowledge of the rates at which calcium
carbonate is lost from the sediments to dissolution.
In practice, it is much easier to determine carbonate
accumulation in the deep sea than it is to estimate
carbonate loss. Yet the latter information is clearly
needed in order to close sediment budgets and to
reconstruct changes in the carbonate system.
Carbonate-rich sediments deposited above the
saturation horizon should experience little in the way
of saturation-driven dissolution because they lie in
contact with waters oversaturated with respect to
calcite. Nevertheless, evidence for significant supralysoclinal dissolution has been found in a number of
studies. Much of this dissolution at shallower water
depths is thought to be driven by chemical reactions
associated with the degradation of organic carbon in
the sediments. Organic carbon arriving at the seafloor is generally respired as CO 2 or remineralized to
other organic compounds by benthic organisms. The
metabolic CO 2 generated by organisms that live
within the sediment can contribute to the dissolution
of calcite even above the lysocline by increasing the
chemical corrosivity of the pore waters. Studies of
organic matter diagenesis in deep-sea sediments
suggest that rates of supralysoclinal dissolution vary
greatly with location, ranging from minimal loss to
440% calcite loss by weight. Temporal and spatial
changes in the rain rate of organic carbon relative to
carbonate can affect this process.
Whether above or below the lysocline, carbonate
dissolution is mostly confined to the bioturbated
surface sediment layer (typically r10 cm in the deep
sea). As carbonate is depleted from this bioturbated
layer, older ‘relict’ carbonate is entrained from the
sediments below. This results in ‘chemical erosion’
and can produce substantial hiatuses or gaps in the
record. Dissolution, and hence erosion, eventually
stops when nonreactive materials fill up the mixed
layer and isolate the underlying sediment from the
overlying water. Many clay layers interbedded within
carbonate-rich sequences are likely produced by this
mechanism; the resulting lithologic contrasts often
show up as subsurface seismic horizons which can be
traced for long distances and tell a story of changing
dissolution gradients and carbonate chemistry in the
past.
Basin-to-basin Fractionation
in the Modern Ocean
Superimposed on the general depth-dependent decrease
of carbonate accumulation observed everywhere in the
deep sea are preservation patterns that differ between
the major ocean basins. Today, carbonate-rich sediments tend to accumulate in the Atlantic Ocean, while
more carbonate-poor sediments are generally found at
comparable water depths in the Indian and Pacific
Oceans. This modern pattern is largely the product of
the ocean’s thermohaline circulation and has been
termed ‘basin-to-basin fractionation’. In the Atlantic,
deep and bottom waters tend to be produced at high
latitudes because cold temperatures and high sea surface salinities lead to the formation of dense water
CALCIUM CARBONATES 341
upper water column while alive. These same characteristics largely dictate their resistance to dissolution
after death. Taxa living in warm, tropical surface
waters, where density is generally low, tend to be
open-structured with thin shells and porous walls.
Taxa that live deeper in cooler, denser subsurface
waters, or in colder surface waters at high latitudes,
tend to be more heavily calcified with thicker shells
and small or closed up pores. On the seafloor, the thinshelled, more fragile species tend to dissolve more
readily than the robust taxa. In effect, this means that
individual species each have their own ‘lysocline’,
which can be offset shallower or deeper from the
foraminiferal lysocline determined from the total assemblage. There are additional consequences of this
selective preservation of taxa that must be considered
in paleo-oceanographic or paleoclimatic studies. For
example, the selective preservation of more heavily
calcified taxa tends to impart a generally ‘cooler’ appearance to the overall microfossil population and
can bias attempts to derive paleotemperature information from seafloor assemblages, as well as other
population properties such as diversity.
For carbonate particles produced in the upper
ocean, settling rates play an important role in their
distribution and preservation. Smaller planktonic
foraminifers settle at about 150–250 m d
À1 , while
larger (4250 mm) foraminifers may settle as much as
2000 m d
À1
. These rates are rapid enough that little
dissolution is thought to occur in the water column.
Solitary coccoliths, on the other hand, sink at rates of
0.3 to B10 m d
À1
, slow enough that dissolution
within the water column should theoretically prevent
their ever reaching the ocean bottom. However,
sediment trap studies have shown that transport by
fecal pellets is the dominant process by which small
phytoplankton skeletons are transferred to the seafloor. Protection offered by the organic fecal pellet
covering may also protect the coccoliths after deposition and account for the fact that the coccolith
lysocline is generally observed to lie somewhat deeper than the foraminiferal lysocline.
While the seafloor depths of the lysocline and
CCD can be readily identified from sedimentary
criteria, this information is of limited use without
realistic knowledge of the rates at which calcium
carbonate is lost from the sediments to dissolution.
In practice, it is much easier to determine carbonate
accumulation in the deep sea than it is to estimate
carbonate loss. Yet the latter information is clearly
needed in order to close sediment budgets and to
reconstruct changes in the carbonate system.
Carbonate-rich sediments deposited above the
saturation horizon should experience little in the way
of saturation-driven dissolution because they lie in
contact with waters oversaturated with respect to
calcite. Nevertheless, evidence for significant supralysoclinal dissolution has been found in a number of
studies. Much of this dissolution at shallower water
depths is thought to be driven by chemical reactions
associated with the degradation of organic carbon in
the sediments. Organic carbon arriving at the seafloor is generally respired as CO 2 or remineralized to
other organic compounds by benthic organisms. The
metabolic CO 2 generated by organisms that live
within the sediment can contribute to the dissolution
of calcite even above the lysocline by increasing the
chemical corrosivity of the pore waters. Studies of
organic matter diagenesis in deep-sea sediments
suggest that rates of supralysoclinal dissolution vary
greatly with location, ranging from minimal loss to
440% calcite loss by weight. Temporal and spatial
changes in the rain rate of organic carbon relative to
carbonate can affect this process.
Whether above or below the lysocline, carbonate
dissolution is mostly confined to the bioturbated
surface sediment layer (typically r10 cm in the deep
sea). As carbonate is depleted from this bioturbated
layer, older ‘relict’ carbonate is entrained from the
sediments below. This results in ‘chemical erosion’
and can produce substantial hiatuses or gaps in the
record. Dissolution, and hence erosion, eventually
stops when nonreactive materials fill up the mixed
layer and isolate the underlying sediment from the
overlying water. Many clay layers interbedded within
carbonate-rich sequences are likely produced by this
mechanism; the resulting lithologic contrasts often
show up as subsurface seismic horizons which can be
traced for long distances and tell a story of changing
dissolution gradients and carbonate chemistry in the
past.
Basin-to-basin Fractionation
in the Modern Ocean
Superimposed on the general depth-dependent decrease
of carbonate accumulation observed everywhere in the
deep sea are preservation patterns that differ between
the major ocean basins. Today, carbonate-rich sediments tend to accumulate in the Atlantic Ocean, while
more carbonate-poor sediments are generally found at
comparable water depths in the Indian and Pacific
Oceans. This modern pattern is largely the product of
the ocean’s thermohaline circulation and has been
termed ‘basin-to-basin fractionation’. In the Atlantic,
deep and bottom waters tend to be produced at high
latitudes because cold temperatures and high sea surface salinities lead to the formation of dense water
CALCIUM CARBONATES 341