in the tropics (Figure 7a). Because the CCD represents the
difference between production and dissolution, high CaCO 3
production forces down the CCD under the equatorial
upwelling region resulting in higher CaCO 3 content in
equatorial sediments (van Andel and Moore, 1974).
Biogenic opal distribution in Figure 7b is presented on
a carbonate-free basis, as a percentage of the noncarbonate
fraction of the sediments. Because biogenic opal is undersaturated in ocean waters, it only accumulates where sedimentation rates are reasonably high. For the most part,
opal is an important but secondary biogenic sediment
component in the pelagic regime. However, diatom oozes
can be found where upwelling brings nutrient-rich waters
to the euphotic zone and high diatom production results.
Diatom oozes are characteristic of the Southern Ocean,
for example (Burckle and Cirilli, 1987). A zone rich in
radiolarian opal has often been noted along the north edge
of the equatorial Pacific upwelling zone, but this primarily
represents the outcropping of Eocene and older biogenic
radiolarian oozes (Riedel, 1967). Eocene radiolarian
oozes result from a shallow CCD prior to 34 million years
ago and conditions favorable to radiolarian production of
biogenic opal. Some researchers have proposed that the
rich deposits of Eocene biogenic opal were caused by elevated dissolved Si in the oceans delivered by a very active
silica weathering cycle, but Moore et al. (2008) showed
that, at least in the equatorial Pacific, biogenic opal burial
was no greater in the Eocene than the Holocene. Additional Si inputs to the oceans are not needed to produce
the ooze deposits.
Both C org and Ba are also biogenic sediment components. Near continental margins, C org sediment contents
are typically 1–2 % and occasionally can reach as high
as ~20 % where hydrodynamically light sediments are
deposited near coastal upwelling zones (Müller and Suess,
1979). In most of the pelagic realm, however, C org contents are less than 0.25 % (Pedersen and Calvert, 1990).
C org burial is enhanced where primary productivity and
sedimentation rates are high and secondarily where bottom water oxygen contents are low (Pedersen and Calvert,
1990). Enhanced C org burial strongly affects the
subseafloor redox environment and causes mobilization
of some elements as well as enhanced burial of
others (section “Seafloor Processes Affecting Deep-Sea
Sediment Composition: Dissolution, Early Diagenesis,
and Sediment Movement”; Froelich et al., 1979;
Calvert and Pedersen, 1993; Morford and Emerson,
1999). Preservation of C org has changed with time, and
C org is more poorly preserved when Earth climates were
warm, perhaps because heterotrophs have higher metabolic demand under warm conditions (Olivarez Lyle and
Lyle, 2006).
Ba is enriched in biogenic sediments as well as in
hydrothermal sediments (Griffith and Paytan, 2012). As
the remains of plankton fall through the water column,
barite (BaSO 4 ) precipitates in microenvironments within
the falling debris (Dehairs et al., 1980; Griffith and Paytan,
2012). Biogenic Ba is the most consistently preserved of
the biogenic sediment components in deep-sea sediments.
In the Holocene, less than 5 % of the C org particulate rain is
typically buried in deep-sea sediments, versus 5–10 % of
biogenic opal, and 30 % biogenic Ba (Dymond et al.,
1992; Dymond and Lyle, 1994). CaCO 3 has highly variable preservation depending on water depth and ocean
chemistry. Where biogenic sediments are severely
dissolved after deposition, a barite (BaSO 4 ) residue can
be the last remaining marker of biogenic deposition
(Dymond, 1981). Barite is preserved in well-oxygenated
sediments but will dissolve under sulfate-reducing conditions and can be remobilized (Torres et al., 2002).
Hydrothermal sediment components
The mid-ocean ridges are a world-circling volcanic mountain chain, creating ~3 km
2 of new hot basalt crust each
year (Williams and von Herzen, 1974). Hot springs that
cool the crust are ubiquitous along mid-ocean ridges, form
rich environments for benthic life (Kelley et al., 2002),
and form distinctive sediments that precipitate around
the vents and are dispersed in hydrothermal plumes.
Unusual sediment composition found on all mid-ocean
ridges led to the hypothesis that hydrothermal systems
must be common in the oceans (Boström and Peterson,
1966, 1969).
Near hydrothermal vents, where large-scale precipitation of minerals from solution results from mixing with
normal ocean water, massive sulfide deposits and oxidized
remains of massive sulfides and other minerals are found
(Haymon, 2005). Proximal hydrothermal deposits can be
highly varied because of the variability of temperature
and mixing conditions. The deposits themselves are
scattered and uncommon because active vent fields are
small relative to the total ridge length and are spaced
~20–50 km apart along the ridge crest (Baker and Urabe,
1996).
Farther away from the actual vents, hydrothermal plumes
consist of particles with much more uniform composition.
The plumes distribute particles rich in amorphous Fe-Mn
oxyhydroxides and poor in Al that form the hydrothermal
sediment component (Boström and Peterson, 1969;
Dymond, 1981). Hydrothermal sediments are most apparent in the South Pacific along the East Pacific Rise, partly
because little besides biogenic carbonates deposit
there and partly because highest ocean crustal formation
occurs there. As the hydrothermal plume material drifts
away from mid-ocean ridges, the Fe-Mn oxyhydroxide particles scavenge other elements from seawater, like P and the
rare earth elements, and form a significant sink for these elements (Wheat et al., 1996; Ruhlin and Owen, 1986).
Downcore variation in the rate that hydrothermal sediments
accumulate has been used to track changes in hydrothermal
activity through time. Lyle et al. (1987) showed that hydrothermal activity at the East Pacific Rise increased dramatically during ridge reorganizations, presumably because
fracturing allowed significantly better seawater access to
hot rock.
DEEP-SEA SEDIMENTS
165
difference between production and dissolution, high CaCO 3
production forces down the CCD under the equatorial
upwelling region resulting in higher CaCO 3 content in
equatorial sediments (van Andel and Moore, 1974).
Biogenic opal distribution in Figure 7b is presented on
a carbonate-free basis, as a percentage of the noncarbonate
fraction of the sediments. Because biogenic opal is undersaturated in ocean waters, it only accumulates where sedimentation rates are reasonably high. For the most part,
opal is an important but secondary biogenic sediment
component in the pelagic regime. However, diatom oozes
can be found where upwelling brings nutrient-rich waters
to the euphotic zone and high diatom production results.
Diatom oozes are characteristic of the Southern Ocean,
for example (Burckle and Cirilli, 1987). A zone rich in
radiolarian opal has often been noted along the north edge
of the equatorial Pacific upwelling zone, but this primarily
represents the outcropping of Eocene and older biogenic
radiolarian oozes (Riedel, 1967). Eocene radiolarian
oozes result from a shallow CCD prior to 34 million years
ago and conditions favorable to radiolarian production of
biogenic opal. Some researchers have proposed that the
rich deposits of Eocene biogenic opal were caused by elevated dissolved Si in the oceans delivered by a very active
silica weathering cycle, but Moore et al. (2008) showed
that, at least in the equatorial Pacific, biogenic opal burial
was no greater in the Eocene than the Holocene. Additional Si inputs to the oceans are not needed to produce
the ooze deposits.
Both C org and Ba are also biogenic sediment components. Near continental margins, C org sediment contents
are typically 1–2 % and occasionally can reach as high
as ~20 % where hydrodynamically light sediments are
deposited near coastal upwelling zones (Müller and Suess,
1979). In most of the pelagic realm, however, C org contents are less than 0.25 % (Pedersen and Calvert, 1990).
C org burial is enhanced where primary productivity and
sedimentation rates are high and secondarily where bottom water oxygen contents are low (Pedersen and Calvert,
1990). Enhanced C org burial strongly affects the
subseafloor redox environment and causes mobilization
of some elements as well as enhanced burial of
others (section “Seafloor Processes Affecting Deep-Sea
Sediment Composition: Dissolution, Early Diagenesis,
and Sediment Movement”; Froelich et al., 1979;
Calvert and Pedersen, 1993; Morford and Emerson,
1999). Preservation of C org has changed with time, and
C org is more poorly preserved when Earth climates were
warm, perhaps because heterotrophs have higher metabolic demand under warm conditions (Olivarez Lyle and
Lyle, 2006).
Ba is enriched in biogenic sediments as well as in
hydrothermal sediments (Griffith and Paytan, 2012). As
the remains of plankton fall through the water column,
barite (BaSO 4 ) precipitates in microenvironments within
the falling debris (Dehairs et al., 1980; Griffith and Paytan,
2012). Biogenic Ba is the most consistently preserved of
the biogenic sediment components in deep-sea sediments.
In the Holocene, less than 5 % of the C org particulate rain is
typically buried in deep-sea sediments, versus 5–10 % of
biogenic opal, and 30 % biogenic Ba (Dymond et al.,
1992; Dymond and Lyle, 1994). CaCO 3 has highly variable preservation depending on water depth and ocean
chemistry. Where biogenic sediments are severely
dissolved after deposition, a barite (BaSO 4 ) residue can
be the last remaining marker of biogenic deposition
(Dymond, 1981). Barite is preserved in well-oxygenated
sediments but will dissolve under sulfate-reducing conditions and can be remobilized (Torres et al., 2002).
Hydrothermal sediment components
The mid-ocean ridges are a world-circling volcanic mountain chain, creating ~3 km
2 of new hot basalt crust each
year (Williams and von Herzen, 1974). Hot springs that
cool the crust are ubiquitous along mid-ocean ridges, form
rich environments for benthic life (Kelley et al., 2002),
and form distinctive sediments that precipitate around
the vents and are dispersed in hydrothermal plumes.
Unusual sediment composition found on all mid-ocean
ridges led to the hypothesis that hydrothermal systems
must be common in the oceans (Boström and Peterson,
1966, 1969).
Near hydrothermal vents, where large-scale precipitation of minerals from solution results from mixing with
normal ocean water, massive sulfide deposits and oxidized
remains of massive sulfides and other minerals are found
(Haymon, 2005). Proximal hydrothermal deposits can be
highly varied because of the variability of temperature
and mixing conditions. The deposits themselves are
scattered and uncommon because active vent fields are
small relative to the total ridge length and are spaced
~20–50 km apart along the ridge crest (Baker and Urabe,
1996).
Farther away from the actual vents, hydrothermal plumes
consist of particles with much more uniform composition.
The plumes distribute particles rich in amorphous Fe-Mn
oxyhydroxides and poor in Al that form the hydrothermal
sediment component (Boström and Peterson, 1969;
Dymond, 1981). Hydrothermal sediments are most apparent in the South Pacific along the East Pacific Rise, partly
because little besides biogenic carbonates deposit
there and partly because highest ocean crustal formation
occurs there. As the hydrothermal plume material drifts
away from mid-ocean ridges, the Fe-Mn oxyhydroxide particles scavenge other elements from seawater, like P and the
rare earth elements, and form a significant sink for these elements (Wheat et al., 1996; Ruhlin and Owen, 1986).
Downcore variation in the rate that hydrothermal sediments
accumulate has been used to track changes in hydrothermal
activity through time. Lyle et al. (1987) showed that hydrothermal activity at the East Pacific Rise increased dramatically during ridge reorganizations, presumably because
fracturing allowed significantly better seawater access to
hot rock.
DEEP-SEA SEDIMENTS
165