9.6 Steady State and Dynamic Denudation Systems
- Scenario 3, constant tectonic uplijt and change 0/
the denudation constant K d (Fig. 9.19a, b and c). Climate change or differences in the erodibility of rocks
lead to changes in Ku (equation 9.2) and thus also in
the response time and "equilibrium elevation" of the
land surface.
Beginning at zero elevation (in relation to the base level of
erosion), continuous uplift first leads to a rising land surface, because denudation lags behind uplift. With increasing elevation, the denudation rate grows linearly or exponentially (Figs. 9.13 and 9.19c) until it matches the rate of
total uplift (tectonic and isostatic rebound). A wet tropical
climate (andJor easily erodable rocks) may bring about
equilibrium conditions earlier and at lower altitude than
temperate humid or arid to semi-arid conditions (Fig.
9.19b). Thus, low denudation rates allow the landscape to
raise to higher plateau or summit elevations, provided tectonic uplift continues.
In all ofthese models, denudation is assumed to proceed
uniformly over the entire area as postulated for the formation of peneplains in tropical to subtropical clirnates. The
influence of different rock types exposed at the same time,
linear downcutting of deep valleys as indicated in Fig.
9.19d, as weIl as other small-scale geomorphological features such as benches, terraces, and pediments, are neglected.
Without denudation, even relatively low uplift rates
as characteristic of epeirogenetic movements (e.g. 20
m1Ma) would generate highly elevated regions in
geological time periods. High uplift rates (e.g. 1000
m1Ma, see scenario 4, below) would produce very
high altitudes within short time periods (10 Ma) if
denudation was absent. However, because the denudation rate grows rapidly with relief, it normally
reaches, with some time lag, astate where it equals
or exceeds the uplift rate of a specific area. The elevation of the land surface above sea level (at present
up to about 9 km) therefore cannot surpass a certain
limit on the Earth's surface. The maximum altitude of
mountain belts or plateaus is also limited by isostasy
depending on rock densities and crustal thicknesses
(cf. Chap. 8.1).
- Scenario 4, orogenic belts (Fig. 9.lge). The model
shows the long-term evolution of the central part of a
fold-thrust belt; it displays a situation similar to that
of large alpino-type mountain belts.
A phase ofstrong uplift (~l mmla = 1000 mlMa) is maintained for about 10 Ma. It is followed by aperiod of comparatively long-term (20 Ma) moderate uplift (ca. 0.3
mmla) caused by isostatic adjustrnent ofthe thickened crost
to previous or decelerating underplating. At the same time,
erosional unloading is compensated to a large extent by
rebound. Both processes control the final stage of the
orogenic system and their rates decrease with time.
At the onset of uplift above base level, the rate of
denudation once again lags behind uplift. After the
453
end of this phase, denudation still proceeds for some
time at a rate almost as high as before. During periods of constant uplift, an equilibrium elevation is
established (at altitudes of ;::3000 m and about 2000
m, respectively) as discussed above (scenario 2). After 50 Ma, arelief of about 500 m is still present.
The two summit levels are, of course, intersected by many
deep valleys, which rnay generate a main valley floor considerably ueeper than the summit elevation (Fig. 9.l8d).
During certain times, for example when uplift surpasses
denudation, channel lowering (D c in Fig. 9.l9d) rnay predominate over summit lowering (D s), ot vice versa. In these
cases, the mean denudation rate is approximately D = 1/2
(Dc + Ds)'
Major. rivers are able to incise through the bedrock at
about the same rate as tectonic uplift. An outstanding example is the upper Indus River in the Nanga Parbat region
where the bedrock was locally raised at a rate of 2-12 mmla
during Quaternary times (Burbank et a1. 1996). The barren
valley slopes maintain a maximum slope angle under these
conditions.
With decreasing uplift and the passage of time, the valleys usually become wider and their slopes gentIer; in this
way the areal extent and heights of the summits are reduced. During the later stages of landscape evolution, when
the summit elevation is already considerably lowered, the
evolving landforms are significantly controlled by clirnatic
factors (Fig. 9.1ge).
In total, rock uplift during the 80 Ma history
amounted to 23 km for this example. This appears to
be realistic, because mountain belts of this age now
exhibit outcropping metamorphie rocks which
formed at burial depths of approximately 10 to 30
km.
There are significant modifications and deviations
from this scenario of the landscape evolution of a
mountain belt. At the front of a foreland thrust-belt,
for example, the rate ofuplift during a short (~5 Ma)
active phase of overthrusting may be so high (up to
10 mmla) that it cannot be compensated for by denudation (e.g. Delcaillau and Carozza 1993; Hurtrez
and Lucazeau 1999; cf. Chap. 12.6.2).
The theoretical background and modelling oflandscapes of
various crigin and scale (for time and space) have been
recently Stüwe 1994; Kooi and Beaumontl996; Pazzaglia and
Brandon 1996; Tucker and Slingerland 1996).
Geomorphologists have become cautious regarding the
usage of the term "equilibrium landforms" (see, e.g.,
Chorley et a1. 1984). It seems that steady-state equilibrium,
with only srnall fluctuations of certain landforms around a
specific average condition, does barely exist in nature. Relatively short-term variations in clirnate and sea-level, especially in the Quaternary, have maintained non-steady conditions in large regions of the world (cf. Chap. 7.9). Landforms resulting from uplift and denudation should therefore
be characterized as dynamic metastable systems. Such systems respond not only to "extemal" influences as discussed
above, but mayaiso be affected by "internai" exogenic pro-
- Scenario 3, constant tectonic uplijt and change 0/
the denudation constant K d (Fig. 9.19a, b and c). Climate change or differences in the erodibility of rocks
lead to changes in Ku (equation 9.2) and thus also in
the response time and "equilibrium elevation" of the
land surface.
Beginning at zero elevation (in relation to the base level of
erosion), continuous uplift first leads to a rising land surface, because denudation lags behind uplift. With increasing elevation, the denudation rate grows linearly or exponentially (Figs. 9.13 and 9.19c) until it matches the rate of
total uplift (tectonic and isostatic rebound). A wet tropical
climate (andJor easily erodable rocks) may bring about
equilibrium conditions earlier and at lower altitude than
temperate humid or arid to semi-arid conditions (Fig.
9.19b). Thus, low denudation rates allow the landscape to
raise to higher plateau or summit elevations, provided tectonic uplift continues.
In all ofthese models, denudation is assumed to proceed
uniformly over the entire area as postulated for the formation of peneplains in tropical to subtropical clirnates. The
influence of different rock types exposed at the same time,
linear downcutting of deep valleys as indicated in Fig.
9.19d, as weIl as other small-scale geomorphological features such as benches, terraces, and pediments, are neglected.
Without denudation, even relatively low uplift rates
as characteristic of epeirogenetic movements (e.g. 20
m1Ma) would generate highly elevated regions in
geological time periods. High uplift rates (e.g. 1000
m1Ma, see scenario 4, below) would produce very
high altitudes within short time periods (10 Ma) if
denudation was absent. However, because the denudation rate grows rapidly with relief, it normally
reaches, with some time lag, astate where it equals
or exceeds the uplift rate of a specific area. The elevation of the land surface above sea level (at present
up to about 9 km) therefore cannot surpass a certain
limit on the Earth's surface. The maximum altitude of
mountain belts or plateaus is also limited by isostasy
depending on rock densities and crustal thicknesses
(cf. Chap. 8.1).
- Scenario 4, orogenic belts (Fig. 9.lge). The model
shows the long-term evolution of the central part of a
fold-thrust belt; it displays a situation similar to that
of large alpino-type mountain belts.
A phase ofstrong uplift (~l mmla = 1000 mlMa) is maintained for about 10 Ma. It is followed by aperiod of comparatively long-term (20 Ma) moderate uplift (ca. 0.3
mmla) caused by isostatic adjustrnent ofthe thickened crost
to previous or decelerating underplating. At the same time,
erosional unloading is compensated to a large extent by
rebound. Both processes control the final stage of the
orogenic system and their rates decrease with time.
At the onset of uplift above base level, the rate of
denudation once again lags behind uplift. After the
453
end of this phase, denudation still proceeds for some
time at a rate almost as high as before. During periods of constant uplift, an equilibrium elevation is
established (at altitudes of ;::3000 m and about 2000
m, respectively) as discussed above (scenario 2). After 50 Ma, arelief of about 500 m is still present.
The two summit levels are, of course, intersected by many
deep valleys, which rnay generate a main valley floor considerably ueeper than the summit elevation (Fig. 9.l8d).
During certain times, for example when uplift surpasses
denudation, channel lowering (D c in Fig. 9.l9d) rnay predominate over summit lowering (D s), ot vice versa. In these
cases, the mean denudation rate is approximately D = 1/2
(Dc + Ds)'
Major. rivers are able to incise through the bedrock at
about the same rate as tectonic uplift. An outstanding example is the upper Indus River in the Nanga Parbat region
where the bedrock was locally raised at a rate of 2-12 mmla
during Quaternary times (Burbank et a1. 1996). The barren
valley slopes maintain a maximum slope angle under these
conditions.
With decreasing uplift and the passage of time, the valleys usually become wider and their slopes gentIer; in this
way the areal extent and heights of the summits are reduced. During the later stages of landscape evolution, when
the summit elevation is already considerably lowered, the
evolving landforms are significantly controlled by clirnatic
factors (Fig. 9.1ge).
In total, rock uplift during the 80 Ma history
amounted to 23 km for this example. This appears to
be realistic, because mountain belts of this age now
exhibit outcropping metamorphie rocks which
formed at burial depths of approximately 10 to 30
km.
There are significant modifications and deviations
from this scenario of the landscape evolution of a
mountain belt. At the front of a foreland thrust-belt,
for example, the rate ofuplift during a short (~5 Ma)
active phase of overthrusting may be so high (up to
10 mmla) that it cannot be compensated for by denudation (e.g. Delcaillau and Carozza 1993; Hurtrez
and Lucazeau 1999; cf. Chap. 12.6.2).
The theoretical background and modelling oflandscapes of
various crigin and scale (for time and space) have been
recently Stüwe 1994; Kooi and Beaumontl996; Pazzaglia and
Brandon 1996; Tucker and Slingerland 1996).
Geomorphologists have become cautious regarding the
usage of the term "equilibrium landforms" (see, e.g.,
Chorley et a1. 1984). It seems that steady-state equilibrium,
with only srnall fluctuations of certain landforms around a
specific average condition, does barely exist in nature. Relatively short-term variations in clirnate and sea-level, especially in the Quaternary, have maintained non-steady conditions in large regions of the world (cf. Chap. 7.9). Landforms resulting from uplift and denudation should therefore
be characterized as dynamic metastable systems. Such systems respond not only to "extemal" influences as discussed
above, but mayaiso be affected by "internai" exogenic pro-