13.4 Early Diagenesis
13.4 Early Diagenesis in
Major Sediment Types
Syndepositional and early diagenesis commonly refers
to changes occurring during burial up to a few hundred
meters, where elevated temperatures are not encountered. The diagenetic system is partially open in this
zone because of exchange of dissolved constituents
with the overlying water body and the underlying strata
(cf. Fig. 13.23a). The processes and products of early
diagenesis vary strongly according to the depositional
environment and the type of sediment.
13.4.1 Marine Muds under Oxic and
Anoxie Conditions
Fine-grained marine sediments deposited slowly under
oxic bottom waters display the following early
diagenetic subzones (Fig. 13.23a,b). The upper sediment layer below the interface is kept in an oxygenated
state by down ward diffusion of dissolved oxygen from
the overlying water column. Hence this layer can be
colonized by benthic organisrns and aerobic bacteria,
which consume and oxidize organic matter reaching
the sea floor. Burrowing organisrns in turn promote
mass transfer from the sea water into the sediment and
vice versa. If all the organic matter is used up and iron
and manganese are oxidized to form oxyhydrates,
brown or reddish brown sediment may result which is
no longer affected by organic processes (cf. Sect.
5.3.3). If some reactive organic matter is left over below the oxidation zone, a very thin zone of bacterial
nitrate reduction and a somewhat thicker zone of anaerobic sulfate reduction develop. In this way, CO2 and
H 2 S (and srnall amounts ofNH 3 ) are produced, which
may migrate to the sea floor by upward diffusion or
react with other ions to form new minerals, for exampie pyrite.
Sediments deposited more rapidly under anoxie
conditions are commonly rich in organic matter and
have astronger tendency to form early diagenetic
authigenic minerals. Nitrate reduction occurs in the
water column and sulfate reduction operates both
above and below the sediment-water interface (Fig.
13.23a,b). Since normal benthic organisms are absent,
comparatively large amounts of organic matter are
available for bacterial sulfate reduction. A shortage in
the sulfate content in the pore water can to some extent
be replenished by downward diffusion of sulfate from
the sea water reservoir. Hence, considerable quantities
of CO 2 and H2S can be produced. Iron released from
minerals and reduced to Fe 2 + is precipitated as
(metastable) monosulfide and stable pyrite, particularly
in the upper part of the sulfate reduction zone. Commonly, early diagenetic pyrite occurs as fine-grained
framboidal grains and aggregates, but it mayaiso form
within internal moulds of fossils. Calcareous fossil
685
shells, particularly those consisting of aragonite, are
frequently dissolved or replaced by pyrite.
Somewhat deeper, but still within the sulfate reduction zone, the decomposition of organic matter causes
the formation of authigenic carbonate, often in the
form of concretions. Diffusive transport of methane
from below may stimulate anaerobic sulfate reduction
and provide the alkalinity for the precipitation of carbonate (CH 4 +S0t= HC0 3 '+HS-+ H 2 S). A slow sedimentation rate or a break in sedimentation favors the
formation of concretions. In particular cases, salinity
variations in the overlying water body may be recorded
in the mineral phases of concretions and their isotopic
composition. Normally, the early formed carbonate
concretions are nonferroan, since the pore fluids are
depleted in iron by the preceding iron sulfide precipitation. Many concretions show a sort of zonation during
their growth which reflects the changing chemical environment with increasing burial depth.
If sulfate is exhausted and can no longer be sufficiently supplied by downward diffusion, the precipitation of pyrite and isotopically light carbonate ceases.
Further decomposition of organic matter is now accomplished by microbial fermentation reactions,
which produce isotopically light methane and heavy
carbon dioxide. Carbon dioxide may be reduced to
provide oxygen for anaerobic respiration. These proces ses are frequently summarized under the term
methanogenesis (Fig. 13.23b). Characteristic authigenic minerals and concretions in this zone are carbonates rich in ferrous iron and l3C (ferroan calcite, dolomite, ankerite, and siderite). In addition, the growth of
iron-rich chlorite minerals has been observed in the
zone of methanogenesis. However, iron-poor marine
muds, for example pelagic carbonates, only form small
quantities of pyrite in the sulfate reduction zone and
little or no authigenic carbonates in both the sulfate
reduction and the methanogenesis zones.
Early iron-bearing carbonates and siderite can form in fi~~grained marine sediments when the rate ofrelease (from SIlicates or iron hydroxide) and reduction of iron exceeds the
rate of sulfate reduction. Then, the dissolved sulfide is insufficient to precipitate all the available dissolved ferrous iron,
as observed in late Holocene marsh and sandflat sediments
on the Norfolk coast, England (Pye et al. 1990). Siderites
formed in such marine environments are commonly impure
(in contrast to freshwater siderites) and rich in Mg and Ca
(Mozley 1989). The concretions tend to develop a flatter
shape than in the sulfate reduction zone. Some ofthe carbonate in the concretions may be derived from dissolution and
reprecipitation of detrital, biogenic carbonate. Even the occurrence of rhodochrosite (MnCO 3) has been reported (e.g.,
von Rad and Botz 1987).
The lowered BC content of carbonate concretions reflects
some bacterial fractionation of carbon isotopes. The lower
boundary of this zone may ]je several meters up to 10m below the sea floor. Early diagenetic phosphogenesis (cf. Sect.
5.3.6) is related to the processes taking place in the sulfate
reduction zone.
13.4 Early Diagenesis in
Major Sediment Types
Syndepositional and early diagenesis commonly refers
to changes occurring during burial up to a few hundred
meters, where elevated temperatures are not encountered. The diagenetic system is partially open in this
zone because of exchange of dissolved constituents
with the overlying water body and the underlying strata
(cf. Fig. 13.23a). The processes and products of early
diagenesis vary strongly according to the depositional
environment and the type of sediment.
13.4.1 Marine Muds under Oxic and
Anoxie Conditions
Fine-grained marine sediments deposited slowly under
oxic bottom waters display the following early
diagenetic subzones (Fig. 13.23a,b). The upper sediment layer below the interface is kept in an oxygenated
state by down ward diffusion of dissolved oxygen from
the overlying water column. Hence this layer can be
colonized by benthic organisrns and aerobic bacteria,
which consume and oxidize organic matter reaching
the sea floor. Burrowing organisrns in turn promote
mass transfer from the sea water into the sediment and
vice versa. If all the organic matter is used up and iron
and manganese are oxidized to form oxyhydrates,
brown or reddish brown sediment may result which is
no longer affected by organic processes (cf. Sect.
5.3.3). If some reactive organic matter is left over below the oxidation zone, a very thin zone of bacterial
nitrate reduction and a somewhat thicker zone of anaerobic sulfate reduction develop. In this way, CO2 and
H 2 S (and srnall amounts ofNH 3 ) are produced, which
may migrate to the sea floor by upward diffusion or
react with other ions to form new minerals, for exampie pyrite.
Sediments deposited more rapidly under anoxie
conditions are commonly rich in organic matter and
have astronger tendency to form early diagenetic
authigenic minerals. Nitrate reduction occurs in the
water column and sulfate reduction operates both
above and below the sediment-water interface (Fig.
13.23a,b). Since normal benthic organisms are absent,
comparatively large amounts of organic matter are
available for bacterial sulfate reduction. A shortage in
the sulfate content in the pore water can to some extent
be replenished by downward diffusion of sulfate from
the sea water reservoir. Hence, considerable quantities
of CO 2 and H2S can be produced. Iron released from
minerals and reduced to Fe 2 + is precipitated as
(metastable) monosulfide and stable pyrite, particularly
in the upper part of the sulfate reduction zone. Commonly, early diagenetic pyrite occurs as fine-grained
framboidal grains and aggregates, but it mayaiso form
within internal moulds of fossils. Calcareous fossil
685
shells, particularly those consisting of aragonite, are
frequently dissolved or replaced by pyrite.
Somewhat deeper, but still within the sulfate reduction zone, the decomposition of organic matter causes
the formation of authigenic carbonate, often in the
form of concretions. Diffusive transport of methane
from below may stimulate anaerobic sulfate reduction
and provide the alkalinity for the precipitation of carbonate (CH 4 +S0t= HC0 3 '+HS-+ H 2 S). A slow sedimentation rate or a break in sedimentation favors the
formation of concretions. In particular cases, salinity
variations in the overlying water body may be recorded
in the mineral phases of concretions and their isotopic
composition. Normally, the early formed carbonate
concretions are nonferroan, since the pore fluids are
depleted in iron by the preceding iron sulfide precipitation. Many concretions show a sort of zonation during
their growth which reflects the changing chemical environment with increasing burial depth.
If sulfate is exhausted and can no longer be sufficiently supplied by downward diffusion, the precipitation of pyrite and isotopically light carbonate ceases.
Further decomposition of organic matter is now accomplished by microbial fermentation reactions,
which produce isotopically light methane and heavy
carbon dioxide. Carbon dioxide may be reduced to
provide oxygen for anaerobic respiration. These proces ses are frequently summarized under the term
methanogenesis (Fig. 13.23b). Characteristic authigenic minerals and concretions in this zone are carbonates rich in ferrous iron and l3C (ferroan calcite, dolomite, ankerite, and siderite). In addition, the growth of
iron-rich chlorite minerals has been observed in the
zone of methanogenesis. However, iron-poor marine
muds, for example pelagic carbonates, only form small
quantities of pyrite in the sulfate reduction zone and
little or no authigenic carbonates in both the sulfate
reduction and the methanogenesis zones.
Early iron-bearing carbonates and siderite can form in fi~~grained marine sediments when the rate ofrelease (from SIlicates or iron hydroxide) and reduction of iron exceeds the
rate of sulfate reduction. Then, the dissolved sulfide is insufficient to precipitate all the available dissolved ferrous iron,
as observed in late Holocene marsh and sandflat sediments
on the Norfolk coast, England (Pye et al. 1990). Siderites
formed in such marine environments are commonly impure
(in contrast to freshwater siderites) and rich in Mg and Ca
(Mozley 1989). The concretions tend to develop a flatter
shape than in the sulfate reduction zone. Some ofthe carbonate in the concretions may be derived from dissolution and
reprecipitation of detrital, biogenic carbonate. Even the occurrence of rhodochrosite (MnCO 3) has been reported (e.g.,
von Rad and Botz 1987).
The lowered BC content of carbonate concretions reflects
some bacterial fractionation of carbon isotopes. The lower
boundary of this zone may ]je several meters up to 10m below the sea floor. Early diagenetic phosphogenesis (cf. Sect.
5.3.6) is related to the processes taking place in the sulfate
reduction zone.