7.9 Cyclo- and Event Stratigraphy
For example, a 13% increase in atmospheric CO 2 has been
postulated to cause a CCD rise ofabout I km (Bender 1984).
In addition, an increase in carbonate production on the
shelves can raise the CCD and Iysocline in the deep sea and
vice versa. Above the Iysocline, carbonate is dissolved only
in the presence of substantial amounts of organic matter (e.g.
Diester-Raass 1991). Furthermore, the extraction of CO 2
from the atmosphere byphytoplankton and expanded vegetation on land ("biological pump") as well as silicate weathering can initiate a cooling trend if carbon is stored in sediments and soils (cf. Sect. 5.6.2). The opposite trend is initiated by release of CO2 into the atmosphere (also see e.g.
Weissert and Mohr 1996).
Lowering of the CCD and improved carbonate preservation in the deep sea can also be promoted by a high river input of dissolved calcium carbonate due to a dense, widely
extended vegetation cover. At the same time, dense vegetation tends to reduce mechanical erosion !lnd the supply of
terrestrial silt and clay to the oceans, thus preventing strong
dilution of carbonate deposition. With the aid of clay minerals, Turonian limestone-rnarl alternations in the Western Interior basin of North America, for example, have been ascribed to wet and dry phases of climate (Sethi and Leithold
1994).
In deep-sea environments variations in carbonate dissolution, due to vertical oscillations of the lysocline,
appear to be the main control on the development of
limestone-marl rhythms. Below upwelling zones,
carbonate-bearing sequences may also show rhythmic
chert accumulations. Black shale deposition and the
generation of redox cycles were favored by a warmer
and more uniform global climate than existing at present (Sect. 7.8).
During times of small temperature differences between the
poles and the equator, such as in the Cretaceous (e.g. Sioan
and Barron 1990), thermohaline oceanic circulation tended
to slow. Instead of cold polar deep-water masses, warm
dense saline water, originating from flooded low-Iatitude
shelves and adjacent seas, may have become the major bottom water source (Arthur et al. 1987; Ray 1996). Since the
solubility of oxygen decreases with increasing temperature
and salinity, the sinking, more saline, warm surface waters
soon became oxygen-depleted. This, in turn, hampered the
mineralization of organic matter in intermediateand deep
waters, as well as the release and recycling of nutrients controlling surface productivity. These processes may have
caused short-term, expanded, intensified oxygen-minimum
zones at intermediate water depths and thus the widespread
deposition of black shale layers, even in epicontinental sea
settings. Subtle changes in climate and oceanic circulation
were sufficient to periodically restore more oxygenated conditions with higher planktonic carbonate production.
A special problem is the identification ofthe various periodicities ofthe Milankovitch frequency band (E, 0, P; Sect.
7.9.2) if, as usual, precise dating ofindividual cycles or thick
packages of a certain kind of cycle is not possible. Then it is
difficult to decide which ofthe different astronomical signals
is recorded by the sedimentary cycles. The situation is better
when P, EI and possibly E2, or 0, EI, and E2 (=410 ka) occur in the same sequence and form 5: I (P:E) or 4: I (E I :E2)
bundles (cf. Fig. 7.1 g). With the aid of such bundles and in
combination with other dating methods, a very high strati365
graphie resolution can be achieved (e.g. Rerbert and D'Rondt
1990; Kauffman et al. 1991; Schwarzacher 1993; Gale
1998).
In conclusion, it appears that the complex feedback
system of the Earth is very sensitive to minor changes
in single processes. The sensitivity of the oceanic circulation system was probably even greater during periods of more uniform global climate than it is today under the conditions of a significant temperature gradient
between the poles and the equator (cf. Sects. 5.2 and
5.6).
7.9.4 Peritidal-Lagoonal and
Subtidal Carbonate Cycles
High-frequency sedimentary cycles, including the
Milankovitch frequency band, have been identified in
siliciclastic, calcareous, and mixed depositional environments. Some examples from lakes have been briefly
discussed in Section 7.7.1, siliciclastics-dominated
parasequences in Sections 7.2.3, 7.3.3, and 7.9.2.
Common and frequently described representatives
of the Milankovitch-type cyclicity are the so-called
peritidal-Iagoonal carbonate cycles which are known
from the Proterozoic up to the Present. These cyclic
sequences usually formed on carbonate platforms including peritidal environments. Their facies frequently
varies between shallow subtidal, intertidal, supratidal,
and subtidal-lagoonal zones which are very sensitive to
even srnall and short changes in sea level or, as an alternative, to laterally migrating sediment bodies.
Biogenie, mostly benthic carbonate production is relatively high in these environments (around 5 to 20
crn/ka, Sect. 10.2) and therefore able to produce
thicker Milankovitch-type depositional cycles (typically one to several meters thick) than planktonic carbonate production generating cyclic bedding in deeper
water. Although peritidal carbonates vary considerably, both laterally and vertically, individual cycles can
be correlated over distances of several km to more than
10 kilometers (Haas 1991), and groups of cycles over
100 km (Elrick 1996; Sect. 7.9.5). Successions of
peritidal-lagoonal cycles commonly reach thicknesses
of several hundreds of meters.
Since the pioneering study by Fischer (1964) in the Triassie
Dachstein limestone ofthe northern Calcareous Alps (known
as "Loferites"), peritidal carbonate cycles have been described from regions all over the world (see, e.g.,
Goldharnmer et al. 1990; Osleger and Read 1991; some articles in Einseie et al. 1991; Pratt et al. 1992; Cowan and
James 1994; Strasser 1994; Elrick 1995; Pasquier and
Strasser 1997; and rnany others). The variation ofperitidal
cycles as a result ofthe interplay between subsidence, sediment aggradation, and composite sea-level changes has been
discussed by several authors (e.g. Koerschner and Read
1989; Goldhammer et al. 1993; Goldhammer and Oswald
1994; Balog et al. 1997).
For example, a 13% increase in atmospheric CO 2 has been
postulated to cause a CCD rise ofabout I km (Bender 1984).
In addition, an increase in carbonate production on the
shelves can raise the CCD and Iysocline in the deep sea and
vice versa. Above the Iysocline, carbonate is dissolved only
in the presence of substantial amounts of organic matter (e.g.
Diester-Raass 1991). Furthermore, the extraction of CO 2
from the atmosphere byphytoplankton and expanded vegetation on land ("biological pump") as well as silicate weathering can initiate a cooling trend if carbon is stored in sediments and soils (cf. Sect. 5.6.2). The opposite trend is initiated by release of CO2 into the atmosphere (also see e.g.
Weissert and Mohr 1996).
Lowering of the CCD and improved carbonate preservation in the deep sea can also be promoted by a high river input of dissolved calcium carbonate due to a dense, widely
extended vegetation cover. At the same time, dense vegetation tends to reduce mechanical erosion !lnd the supply of
terrestrial silt and clay to the oceans, thus preventing strong
dilution of carbonate deposition. With the aid of clay minerals, Turonian limestone-rnarl alternations in the Western Interior basin of North America, for example, have been ascribed to wet and dry phases of climate (Sethi and Leithold
1994).
In deep-sea environments variations in carbonate dissolution, due to vertical oscillations of the lysocline,
appear to be the main control on the development of
limestone-marl rhythms. Below upwelling zones,
carbonate-bearing sequences may also show rhythmic
chert accumulations. Black shale deposition and the
generation of redox cycles were favored by a warmer
and more uniform global climate than existing at present (Sect. 7.8).
During times of small temperature differences between the
poles and the equator, such as in the Cretaceous (e.g. Sioan
and Barron 1990), thermohaline oceanic circulation tended
to slow. Instead of cold polar deep-water masses, warm
dense saline water, originating from flooded low-Iatitude
shelves and adjacent seas, may have become the major bottom water source (Arthur et al. 1987; Ray 1996). Since the
solubility of oxygen decreases with increasing temperature
and salinity, the sinking, more saline, warm surface waters
soon became oxygen-depleted. This, in turn, hampered the
mineralization of organic matter in intermediateand deep
waters, as well as the release and recycling of nutrients controlling surface productivity. These processes may have
caused short-term, expanded, intensified oxygen-minimum
zones at intermediate water depths and thus the widespread
deposition of black shale layers, even in epicontinental sea
settings. Subtle changes in climate and oceanic circulation
were sufficient to periodically restore more oxygenated conditions with higher planktonic carbonate production.
A special problem is the identification ofthe various periodicities ofthe Milankovitch frequency band (E, 0, P; Sect.
7.9.2) if, as usual, precise dating ofindividual cycles or thick
packages of a certain kind of cycle is not possible. Then it is
difficult to decide which ofthe different astronomical signals
is recorded by the sedimentary cycles. The situation is better
when P, EI and possibly E2, or 0, EI, and E2 (=410 ka) occur in the same sequence and form 5: I (P:E) or 4: I (E I :E2)
bundles (cf. Fig. 7.1 g). With the aid of such bundles and in
combination with other dating methods, a very high strati365
graphie resolution can be achieved (e.g. Rerbert and D'Rondt
1990; Kauffman et al. 1991; Schwarzacher 1993; Gale
1998).
In conclusion, it appears that the complex feedback
system of the Earth is very sensitive to minor changes
in single processes. The sensitivity of the oceanic circulation system was probably even greater during periods of more uniform global climate than it is today under the conditions of a significant temperature gradient
between the poles and the equator (cf. Sects. 5.2 and
5.6).
7.9.4 Peritidal-Lagoonal and
Subtidal Carbonate Cycles
High-frequency sedimentary cycles, including the
Milankovitch frequency band, have been identified in
siliciclastic, calcareous, and mixed depositional environments. Some examples from lakes have been briefly
discussed in Section 7.7.1, siliciclastics-dominated
parasequences in Sections 7.2.3, 7.3.3, and 7.9.2.
Common and frequently described representatives
of the Milankovitch-type cyclicity are the so-called
peritidal-Iagoonal carbonate cycles which are known
from the Proterozoic up to the Present. These cyclic
sequences usually formed on carbonate platforms including peritidal environments. Their facies frequently
varies between shallow subtidal, intertidal, supratidal,
and subtidal-lagoonal zones which are very sensitive to
even srnall and short changes in sea level or, as an alternative, to laterally migrating sediment bodies.
Biogenie, mostly benthic carbonate production is relatively high in these environments (around 5 to 20
crn/ka, Sect. 10.2) and therefore able to produce
thicker Milankovitch-type depositional cycles (typically one to several meters thick) than planktonic carbonate production generating cyclic bedding in deeper
water. Although peritidal carbonates vary considerably, both laterally and vertically, individual cycles can
be correlated over distances of several km to more than
10 kilometers (Haas 1991), and groups of cycles over
100 km (Elrick 1996; Sect. 7.9.5). Successions of
peritidal-lagoonal cycles commonly reach thicknesses
of several hundreds of meters.
Since the pioneering study by Fischer (1964) in the Triassie
Dachstein limestone ofthe northern Calcareous Alps (known
as "Loferites"), peritidal carbonate cycles have been described from regions all over the world (see, e.g.,
Goldharnmer et al. 1990; Osleger and Read 1991; some articles in Einseie et al. 1991; Pratt et al. 1992; Cowan and
James 1994; Strasser 1994; Elrick 1995; Pasquier and
Strasser 1997; and rnany others). The variation ofperitidal
cycles as a result ofthe interplay between subsidence, sediment aggradation, and composite sea-level changes has been
discussed by several authors (e.g. Koerschner and Read
1989; Goldhammer et al. 1993; Goldhammer and Oswald
1994; Balog et al. 1997).