6.3 Red Beds
primary nature of these red beds is, however, not
available. They mayaiso result from originally
brown material or reflect post-depositional oxidation
processes (see below).
In Situ (Secondary) Formation of Red Beds
Many red beds do not result from redeposited
lateritic soils, but reflect in situ diagenetic processes
and may therefore be referred to as secondary red
beds. In this case, the source area and its climate are
of minor importance. The principal factor controlling
the subsequent sediment color is a weil oxygenated
depositional environment in which the organic matter, either autochthonous or allochthonous, is mineralized completely or to a very high degree. Then the
potential of the relic organic matter for reducing all
ferric iron compounds present in the sediments is too
low (Fig. 6.3).
The precursors of such red beds are either brown
soils or brown sediments, because ferric iron
hydroxides left behind after the consumption of organic matter are commonly brown or ocher in color
at or near the sediment surface. With increasing
burial depth, temperature, and age, these iron compounds lose their water and are transformed into hematite. Hence, one can generally expect that presentday brown sediments, found below the zone of active
aeration and bottom life, alter their color after burial
and become red. The formation and preservation of
brownJred colored sediments are favored by:
- High iron content in the sediment (e.g., siliciclastic
material rich in biotite, pyroxene, amphibole, olivine,
and iron-bearing clay minerals).
- Low production and rapid mineralization of organic matter in the depositional area (e.g., fluvial
plain with sparse vegetation in semiarid climate).
- Slow or intermittent sedimentation, allowing the
sparse organic matter to be largely consumed in the
soil zone or at the floor of lakes and the sea.
- Redox and pH conditions in interstitial water favoring the formation of iron oxide after burial, if sub sequent reduction is absent.
Depositional Environments
of Brown and Red Beds
Examples of depositional environments with a high
potential for forming brown beds can be easily deduced from these general roles. They include both
continental and marine systems (Fig. 6.3b). The fluvial red muds tones of many formations can be interpreted as originally brown floodplain sediments derived from muddy suspension load or aggregated
river bed load (Rust and Nanson 1989; cf. Sect.
2.2.3).
257
Environments with relatively rapid sedimentation
such· as many river floodplains, prodelta deposits in
lakes and in the sea, or supratidal sediments can
maintain an oxidized state only when their primary
organic carbon content is low. In contrast, slowly
deposited pelagic sediments on submarine plateaus
and in the deep sea may originally contain relatively
high amounts of organic carbon, but benthic life in
oxygenated waters has sufficient time to destroy the
organic :natter before it is finally buried under younger sediments. Red deep-sea clay accumulating below the calcite compensation depth (Sect. 5.3.2) is
representative of this group of red beds, which does
not depend on a specific clay mineral source.
Red deep-sea sediments are restricted to basinal sites and
mid-oceanic ridges without major clastic influx (Franke
and Paul 1980). They rnay contain biogenic silica and carbonate. Red clays sometimes alternate with gray turbidites
which are deposited rapidly and obtain their organic material from source areas with higher sedimentation rates
(Faupl and Sauer 1978).
Carbonates poor in iron only require a little preserved organic carbon to be held in a reduced state.
Therefore, most ancient carbonates are light gray in
color. Red limestones indicate a particularly well oxidized environment, a low sedimentation rate, and in
some cases the emergence above sea level.
Some red limestones, sandstones, and to a lesser
degree claystones with irregular color boundaries
across primary bedding planes may result from
postdepositional processes, such as circulating
groundwaters with a high redox potential. Similarly,
reducing pore waters can destroy red staining, partic~
ularly along zones of high fracture permeability (Fig.
6.3e).
Transgressions ofthe sea over brown or red continental s{;diments normally leave behind gray deposits, but can also affect the underlying sediments by
providing circulating, reducing pore waters. Thus,
the uppermost portions and locally deeper zones of
these sediments may be discolored.
Finally, it is important to note that iron is also released from unstable, iron-bearing minerals during
diagenesis. Oxidizing interstitial waters lead to in
situ weathering of these particles and to the formation of new clay minerals, as well as ferric iron hydroxide. The latter slowly alters to hematite and thus
causes red staining as mentioned earlier. This process
commonly takes place in sediments which were originally brown and did not contain sufficient organic
matter to maintain reduced conditions in the
subsurface.
primary nature of these red beds is, however, not
available. They mayaiso result from originally
brown material or reflect post-depositional oxidation
processes (see below).
In Situ (Secondary) Formation of Red Beds
Many red beds do not result from redeposited
lateritic soils, but reflect in situ diagenetic processes
and may therefore be referred to as secondary red
beds. In this case, the source area and its climate are
of minor importance. The principal factor controlling
the subsequent sediment color is a weil oxygenated
depositional environment in which the organic matter, either autochthonous or allochthonous, is mineralized completely or to a very high degree. Then the
potential of the relic organic matter for reducing all
ferric iron compounds present in the sediments is too
low (Fig. 6.3).
The precursors of such red beds are either brown
soils or brown sediments, because ferric iron
hydroxides left behind after the consumption of organic matter are commonly brown or ocher in color
at or near the sediment surface. With increasing
burial depth, temperature, and age, these iron compounds lose their water and are transformed into hematite. Hence, one can generally expect that presentday brown sediments, found below the zone of active
aeration and bottom life, alter their color after burial
and become red. The formation and preservation of
brownJred colored sediments are favored by:
- High iron content in the sediment (e.g., siliciclastic
material rich in biotite, pyroxene, amphibole, olivine,
and iron-bearing clay minerals).
- Low production and rapid mineralization of organic matter in the depositional area (e.g., fluvial
plain with sparse vegetation in semiarid climate).
- Slow or intermittent sedimentation, allowing the
sparse organic matter to be largely consumed in the
soil zone or at the floor of lakes and the sea.
- Redox and pH conditions in interstitial water favoring the formation of iron oxide after burial, if sub sequent reduction is absent.
Depositional Environments
of Brown and Red Beds
Examples of depositional environments with a high
potential for forming brown beds can be easily deduced from these general roles. They include both
continental and marine systems (Fig. 6.3b). The fluvial red muds tones of many formations can be interpreted as originally brown floodplain sediments derived from muddy suspension load or aggregated
river bed load (Rust and Nanson 1989; cf. Sect.
2.2.3).
257
Environments with relatively rapid sedimentation
such· as many river floodplains, prodelta deposits in
lakes and in the sea, or supratidal sediments can
maintain an oxidized state only when their primary
organic carbon content is low. In contrast, slowly
deposited pelagic sediments on submarine plateaus
and in the deep sea may originally contain relatively
high amounts of organic carbon, but benthic life in
oxygenated waters has sufficient time to destroy the
organic :natter before it is finally buried under younger sediments. Red deep-sea clay accumulating below the calcite compensation depth (Sect. 5.3.2) is
representative of this group of red beds, which does
not depend on a specific clay mineral source.
Red deep-sea sediments are restricted to basinal sites and
mid-oceanic ridges without major clastic influx (Franke
and Paul 1980). They rnay contain biogenic silica and carbonate. Red clays sometimes alternate with gray turbidites
which are deposited rapidly and obtain their organic material from source areas with higher sedimentation rates
(Faupl and Sauer 1978).
Carbonates poor in iron only require a little preserved organic carbon to be held in a reduced state.
Therefore, most ancient carbonates are light gray in
color. Red limestones indicate a particularly well oxidized environment, a low sedimentation rate, and in
some cases the emergence above sea level.
Some red limestones, sandstones, and to a lesser
degree claystones with irregular color boundaries
across primary bedding planes may result from
postdepositional processes, such as circulating
groundwaters with a high redox potential. Similarly,
reducing pore waters can destroy red staining, partic~
ularly along zones of high fracture permeability (Fig.
6.3e).
Transgressions ofthe sea over brown or red continental s{;diments normally leave behind gray deposits, but can also affect the underlying sediments by
providing circulating, reducing pore waters. Thus,
the uppermost portions and locally deeper zones of
these sediments may be discolored.
Finally, it is important to note that iron is also released from unstable, iron-bearing minerals during
diagenesis. Oxidizing interstitial waters lead to in
situ weathering of these particles and to the formation of new clay minerals, as well as ferric iron hydroxide. The latter slowly alters to hematite and thus
causes red staining as mentioned earlier. This process
commonly takes place in sediments which were originally brown and did not contain sufficient organic
matter to maintain reduced conditions in the
subsurface.