11.3 Dynamic Denudation Accumulation Systems
terrigenous deposits (Van Andel et al. 1977). With continued sea-floor spreading in the late Mesozoic and early Cenozoic, the sedimentation rate of clastics dropped to values
mostly under 10 mlMa along the margins and to 2 to 5
mlMa on the Mid-Atlantic ridge (Fig. 11.15b). In contrast,
the in situ biogenic carbonate production and preservation
did not show such a distinct trend, but rather was controlled by changes in the oceanic circullition system.
North Atlantic. Similar to the South Atlantic, the
North Atlantic shows sedimentation rates of
terrigenous material which were generally high during its early history and then declined with the widening of the basin up to the early Miocene (Ehrmann
and Thiede 1985). However, this rate increased again
from the middle Miocene up to. the Quatemary as a
result of climatic change and glacial action in high
latitudes regions and high mountain ranges.
An approximate mass balance of the present-day denudation on land and overall sedimentation in the Atlantic
Ocean between 50 0 N and S yields the following results.
The total drainage area in North and South America, Europe, and Africa is 30 x 10 6 km 2 and the basin area 75 x 10 6
km 2 • Hence the AjAb ratio is 0040. Assuming the modem
global mechanical denudation rate of 50 mlMa (Sect. 9.3)
for the entire drainage area, the average sedimentation rate
of clastic material in the Atlantic is about 50 x 004 = 20
mlMa (solid, water-free material). Compared with Figure
11.15b, this value appears to be realistic. However, this
estimate does not take into account that much sediment is
deposited on the shelves ofthis ocean basin.
As regards chemical denudation, the global rate is 16
m1Ma, and approximately 12% of the dissolved river
load is Ca 2 + and 10% is Si0 2 • Calcium denudation of
~2 m1Ma may lead to a calcium carbonate sedimentation rate (factor 2.5) in the Atlantic of 2 x 2.5 x 0.4 =
2 mlMa or 2 mmlka (AiA b = 0.4). This value is
lower than the actual rate observed above the CCD,
but higher than that found below the CCD where carbonate is largely dissolved. The present-day river
supply of calcium is insufficient to maintain the high
oceanic biogenic carbonate production. The deficiency of the ocean in calcium causes carbonate dissolution below the CCD (cf. Sect. 5.3.2). Similarly,
the river-bome Si0 2 enables an average sedimentation rate of 1.6 mlMa or 1.6 mmlka. This value is too
low to provide enough silica for the ongoing production of siliceous skeletons by diatoms, radiolaria, and
sponges. Hence, most oftheir skeletons are dissolved
in the water column (Sect. 5.3.5).
11.3.4 Closing Basin Model
The most important group of closing basins are related to subduction and plate collision. The general
principles introduced for extensional basins can also
be applied to closing basins which therefore are discussed only briefly.
505
For the model of a closing basin (Fig. 11.16) both
sediment supply and the size of the drainage area, A d ,
are assumed to be constant as for the previous models of extensional basins (Fig. l3a). Subsidence follows the rule found for aging, cooling oceanic crust
and is enforced by the increasing sediment load.
With Ad= constant, the AiAb ratio of the closing
ocean basin increases rapidly once it has become already narrow (remnant basins, Fig. 11.l3a). Then the
mean mechanical sedimentation rate and thus the
accumulating sediment thickness increase exponentially with time and lead to rapid filling of the initially deep basin (Fig. 11.l6a and b).
Initially, the remnant basin is deep and its rates of sedimentation and subsidence, controlled mainly by cooling of the
underlying oceanic crust, are slow. Approaching the
subduction zone, the oceanic crust tends to be flexed, leading to increased subsidence of the basin floor. Simultaneously, the sedimentation rate grows and thus compensates for subsidence. Finally, the sedimentation rate greatly
surpasses subsidence causing rapid and complete basin
filling. This can be frequently observed in orogenic belts
where flysch sequences pass upward into shallow-water
and continental deposits. As a result of subduction and the
increasing sediment load, the top of the oceanic crust may
become deeply buried. This is also known from passive
continental margins below prograding marine deltas (e.g.
the Niger delta) or below the head of large deep-sea fans
(e.g. the Bengal fan, Sect. 11.5.5). Uplift of the magmatic
arc and an increasing sediment supply from the arc region
would further accelerate basin filling.
11.3.5 Foreland Basin Model
Foreland basins in front of overthrust belts (Fig.
11.17a; cf. Sect. 12.6) are in many ways more complex than the previous examples of basin evolution.
The elementary model of Fig. 11.17 shows a first,
pronounced phase of overthrusting and the resulting
increase in erosion and basin filling.
The model is based on the following assumptions: (1) the
size or width ofthe basin area and its drainage area remain
constant (AjAb=I). The mean mechanical sedimentation
rate, SR",e (here for compact sediment), in the basin is
therefore equal to the mean mechanical denudation rate,
DR.ne, in the drainage area. (2) SR me increases from 10 to
250 mmlka during an active phase of overthrusting and
decreases thereafter. (3) The subsidence history of a certain
location in the basin resembles that observed in actual foreland basins. It is mainly controlled by flexural downwarping of continental crust in response to the loads of the
overthrust sheets and the accumulating sediments (cf.
Chap. 8). (4) After overthrusting has ceased, erosion continues at a reduced rate causing so me rebound (uplift) of
the denudational area. This also affects to some extent the
basin area.
Prior to the approach of the overthrust belt, the basin
is shallow-marine (or continental), slowly subsiding
terrigenous deposits (Van Andel et al. 1977). With continued sea-floor spreading in the late Mesozoic and early Cenozoic, the sedimentation rate of clastics dropped to values
mostly under 10 mlMa along the margins and to 2 to 5
mlMa on the Mid-Atlantic ridge (Fig. 11.15b). In contrast,
the in situ biogenic carbonate production and preservation
did not show such a distinct trend, but rather was controlled by changes in the oceanic circullition system.
North Atlantic. Similar to the South Atlantic, the
North Atlantic shows sedimentation rates of
terrigenous material which were generally high during its early history and then declined with the widening of the basin up to the early Miocene (Ehrmann
and Thiede 1985). However, this rate increased again
from the middle Miocene up to. the Quatemary as a
result of climatic change and glacial action in high
latitudes regions and high mountain ranges.
An approximate mass balance of the present-day denudation on land and overall sedimentation in the Atlantic
Ocean between 50 0 N and S yields the following results.
The total drainage area in North and South America, Europe, and Africa is 30 x 10 6 km 2 and the basin area 75 x 10 6
km 2 • Hence the AjAb ratio is 0040. Assuming the modem
global mechanical denudation rate of 50 mlMa (Sect. 9.3)
for the entire drainage area, the average sedimentation rate
of clastic material in the Atlantic is about 50 x 004 = 20
mlMa (solid, water-free material). Compared with Figure
11.15b, this value appears to be realistic. However, this
estimate does not take into account that much sediment is
deposited on the shelves ofthis ocean basin.
As regards chemical denudation, the global rate is 16
m1Ma, and approximately 12% of the dissolved river
load is Ca 2 + and 10% is Si0 2 • Calcium denudation of
~2 m1Ma may lead to a calcium carbonate sedimentation rate (factor 2.5) in the Atlantic of 2 x 2.5 x 0.4 =
2 mlMa or 2 mmlka (AiA b = 0.4). This value is
lower than the actual rate observed above the CCD,
but higher than that found below the CCD where carbonate is largely dissolved. The present-day river
supply of calcium is insufficient to maintain the high
oceanic biogenic carbonate production. The deficiency of the ocean in calcium causes carbonate dissolution below the CCD (cf. Sect. 5.3.2). Similarly,
the river-bome Si0 2 enables an average sedimentation rate of 1.6 mlMa or 1.6 mmlka. This value is too
low to provide enough silica for the ongoing production of siliceous skeletons by diatoms, radiolaria, and
sponges. Hence, most oftheir skeletons are dissolved
in the water column (Sect. 5.3.5).
11.3.4 Closing Basin Model
The most important group of closing basins are related to subduction and plate collision. The general
principles introduced for extensional basins can also
be applied to closing basins which therefore are discussed only briefly.
505
For the model of a closing basin (Fig. 11.16) both
sediment supply and the size of the drainage area, A d ,
are assumed to be constant as for the previous models of extensional basins (Fig. l3a). Subsidence follows the rule found for aging, cooling oceanic crust
and is enforced by the increasing sediment load.
With Ad= constant, the AiAb ratio of the closing
ocean basin increases rapidly once it has become already narrow (remnant basins, Fig. 11.l3a). Then the
mean mechanical sedimentation rate and thus the
accumulating sediment thickness increase exponentially with time and lead to rapid filling of the initially deep basin (Fig. 11.l6a and b).
Initially, the remnant basin is deep and its rates of sedimentation and subsidence, controlled mainly by cooling of the
underlying oceanic crust, are slow. Approaching the
subduction zone, the oceanic crust tends to be flexed, leading to increased subsidence of the basin floor. Simultaneously, the sedimentation rate grows and thus compensates for subsidence. Finally, the sedimentation rate greatly
surpasses subsidence causing rapid and complete basin
filling. This can be frequently observed in orogenic belts
where flysch sequences pass upward into shallow-water
and continental deposits. As a result of subduction and the
increasing sediment load, the top of the oceanic crust may
become deeply buried. This is also known from passive
continental margins below prograding marine deltas (e.g.
the Niger delta) or below the head of large deep-sea fans
(e.g. the Bengal fan, Sect. 11.5.5). Uplift of the magmatic
arc and an increasing sediment supply from the arc region
would further accelerate basin filling.
11.3.5 Foreland Basin Model
Foreland basins in front of overthrust belts (Fig.
11.17a; cf. Sect. 12.6) are in many ways more complex than the previous examples of basin evolution.
The elementary model of Fig. 11.17 shows a first,
pronounced phase of overthrusting and the resulting
increase in erosion and basin filling.
The model is based on the following assumptions: (1) the
size or width ofthe basin area and its drainage area remain
constant (AjAb=I). The mean mechanical sedimentation
rate, SR",e (here for compact sediment), in the basin is
therefore equal to the mean mechanical denudation rate,
DR.ne, in the drainage area. (2) SR me increases from 10 to
250 mmlka during an active phase of overthrusting and
decreases thereafter. (3) The subsidence history of a certain
location in the basin resembles that observed in actual foreland basins. It is mainly controlled by flexural downwarping of continental crust in response to the loads of the
overthrust sheets and the accumulating sediments (cf.
Chap. 8). (4) After overthrusting has ceased, erosion continues at a reduced rate causing so me rebound (uplift) of
the denudational area. This also affects to some extent the
basin area.
Prior to the approach of the overthrust belt, the basin
is shallow-marine (or continental), slowly subsiding