9.6 Steady State and Dynamic Denudation Systems
diocarbon dates of coral reefs raised above sea level during
the Holocene (Dorsey 1989). Along the Cascadia
subduction zone on the northwestern coast of the United
States, young surface uplift ranges from 0 to 5 mrnIa
(Mitchell et al. 1994). Surface up1ift in Japan, associated
with an active subduction zone, amounts to 1 to 4 mrnIa.
Parts of the North Island of New Zea1and show young,
subduction-related surface up1ift of > 1 mrnIa, locally >3
mrnIa (Walcott 1987). For the Caucasus, surface uplift
rates of about 10 mrnIa have been reported (International
Union ofGeodesy and Geophysics 1975).
Uplift of orogenie hefts is a lengthy process lasting
between 20 and 100 Ma Tectonic uplift generally
evolves as a result of plate convergence, subduction,
and underthrusting of oceanic or continental crust.
Later, isostatic rebound of the thickened mountain
root is the principal factor of slowly declining uplift
rates. In detail, the history of mountain ranges is often more complex. Crustal thickening and updorning
of central zones frequently causes detachment zones
and backthrusting, i.e. the upper crust and its sediment cover move, driven by gravity forces, away
from the centers of tectonic uplift. This process is
called "tectonic exhumation" and does not provide
sediment. Furthermore, the structures of orogenic
belts or parts of them may be significantly modified
by lateral extrusion tectonics (e.g. Ratschbacher et al.
1991; Frisch et al. 1998). As a consequence, crustal
blocks (or basins) separated by strike-slip motions,
may experience different histories of tectonic uplift
and exhumation. Average rates of uplift and denudation for entire orogenic belts are therefore problematic.
Exhumation of the Tauern Window, eastern Alps, mainly
occurred in the Miocene at a rate of 1-2 mrnIa which was
largely due to tectonic exhumation (Hey11997; Frisch et al.
1998). Average denudation rates for individual parts ofthis
mountain range and longer time periods range from 0.02 to
0.5 mrnIa. The western (Swiss) Alps mainly evolved under
the influence of north-south convergence and gravity tectonics, but were less affected by lateral extrusion than the
eastern Alps (Schmid et al. 1996). Here, exhumation also
varied significantly in space and time reaching peak values
of up to 5 mrnIa in some crystalline massifs (e.g. the
Bergell) for specific time periods (5-10 Ma). Long-term
denudation rates again are much slower (on the order ofO.5
mrnIa in the crystalline massifs and <0.1 mrnIa in the nappe
zones where sedimentary rocks are exposed. In the Pyrenean orogen, uplift-driven denudation in the time period
from 40 to 10 Ma B.P. mostly operated at rates of 35-60
mrnIka (max. 240 mmlka) but varied spatially and temporally (Morris et al. 1998).
The uplift history of the Tibetan Plateau seems to be
controversial. The long-term (25 Ma) average uplift rate
may have been 200 to 300 mmJka (Zhao and Morgan
1985). The High Himalayas reached an average rate of
1000 to 1500 mmJka during the past 20 Ma (Burg et al.
1987; Einsele et al. 1996), but locally and temporarily the
rates varied. The uplift rate in the Nanga-Parbat-Haramosh
massif, for example, increased from :;0.5 to more than 10
451
mm1a during the past 7 Ma (Hurford 1991; Burbank et al.
1996). In the backthrust zone of the Transhimalayas with
exposed sedimentary rocks, the long-term (20 Ma) average
rate of denudation was 100-200 mmlka only (cf. Sect.
11.5.5).
Modern rates of tectonic uplift are sometimes unusually high because of delayed isostatic rebound in response to ice melted after the last glacial period.
Evaluating repeated geodetic surveys or dated elevated fluvial terraces, very high uplift rates (~3
mm1a) were found in several high mountain ranges
(Gansser 1984). Using geomorphological marker horizons, Darmendrail et al. (1993) found uplift rates of
1 to 3 mrnIa for the past 100 ka in the French Alps.
In the less glaciated Umbrian Apennines, central
Italy, uplift rates in the Quatemary ranged from ~250
to 500 mmlka (Rasse 1993). More data ofuplift rates
in various tectonic settings are summarized in special
articles (e.g. Leeder 1991).
9.6.3 Simple Uplift-Denudation Models
Substantial denudation is commonly initiated by tectonic and· thermal uplift (both included in the term
"tectonic" uplift in the following). Later, after the
end of this process, rock uplift is maintained to some
extent by isostatic rebound following erosional unloading as mentioned above (Fig. 9. 17a). Repeated or
changing vertical tectonic motions and their effects
on denudation and landscape evolution are demonstrated in the models shown in Fig. 9.18a and b and
Fig.9.151:
- Scenario 1, teetonie pulses (Fig. 9.18a). Lowland
(around base level) is raised rapidly by tectonic uplift
during a short time period. After this event, landscape lowering is controlled by denudation and isostatic rebound until a second tectonic pulse increases
mean elevation and thus generates a second period of
slow, exponential decay of the land surface. In this
case, no equilibrium between the (changing) denudation rate and total uplift during the pulses is reached
because the response time of the system is longer
than the tectonic pulse ("Penck-type" landscape evolution; see e.g. Kooi and Beaumont 1996). Finally, a
short period of subsidence brings about accelerated
surface lowering followed again by slow decay.
- Scenario 2, eontinuous lang-term upliji (Fig.
9.18b). Rapidly changing rates of epeirogenic uplift
force denudation to follow until, after some time, a
kind of dynamic equilibrium (steady-state condition)
between total uplift and denudation is obtained
("Hack-type" landscapes). After each change in the
rate of uplift, the system approaches a new dynamic
equilibrium at higher or lower elevation above base
level. The equilibrium level also follows a lowered
base level with a certain response time, t c •
diocarbon dates of coral reefs raised above sea level during
the Holocene (Dorsey 1989). Along the Cascadia
subduction zone on the northwestern coast of the United
States, young surface uplift ranges from 0 to 5 mrnIa
(Mitchell et al. 1994). Surface up1ift in Japan, associated
with an active subduction zone, amounts to 1 to 4 mrnIa.
Parts of the North Island of New Zea1and show young,
subduction-related surface up1ift of > 1 mrnIa, locally >3
mrnIa (Walcott 1987). For the Caucasus, surface uplift
rates of about 10 mrnIa have been reported (International
Union ofGeodesy and Geophysics 1975).
Uplift of orogenie hefts is a lengthy process lasting
between 20 and 100 Ma Tectonic uplift generally
evolves as a result of plate convergence, subduction,
and underthrusting of oceanic or continental crust.
Later, isostatic rebound of the thickened mountain
root is the principal factor of slowly declining uplift
rates. In detail, the history of mountain ranges is often more complex. Crustal thickening and updorning
of central zones frequently causes detachment zones
and backthrusting, i.e. the upper crust and its sediment cover move, driven by gravity forces, away
from the centers of tectonic uplift. This process is
called "tectonic exhumation" and does not provide
sediment. Furthermore, the structures of orogenic
belts or parts of them may be significantly modified
by lateral extrusion tectonics (e.g. Ratschbacher et al.
1991; Frisch et al. 1998). As a consequence, crustal
blocks (or basins) separated by strike-slip motions,
may experience different histories of tectonic uplift
and exhumation. Average rates of uplift and denudation for entire orogenic belts are therefore problematic.
Exhumation of the Tauern Window, eastern Alps, mainly
occurred in the Miocene at a rate of 1-2 mrnIa which was
largely due to tectonic exhumation (Hey11997; Frisch et al.
1998). Average denudation rates for individual parts ofthis
mountain range and longer time periods range from 0.02 to
0.5 mrnIa. The western (Swiss) Alps mainly evolved under
the influence of north-south convergence and gravity tectonics, but were less affected by lateral extrusion than the
eastern Alps (Schmid et al. 1996). Here, exhumation also
varied significantly in space and time reaching peak values
of up to 5 mrnIa in some crystalline massifs (e.g. the
Bergell) for specific time periods (5-10 Ma). Long-term
denudation rates again are much slower (on the order ofO.5
mrnIa in the crystalline massifs and <0.1 mrnIa in the nappe
zones where sedimentary rocks are exposed. In the Pyrenean orogen, uplift-driven denudation in the time period
from 40 to 10 Ma B.P. mostly operated at rates of 35-60
mrnIka (max. 240 mmlka) but varied spatially and temporally (Morris et al. 1998).
The uplift history of the Tibetan Plateau seems to be
controversial. The long-term (25 Ma) average uplift rate
may have been 200 to 300 mmJka (Zhao and Morgan
1985). The High Himalayas reached an average rate of
1000 to 1500 mmJka during the past 20 Ma (Burg et al.
1987; Einsele et al. 1996), but locally and temporarily the
rates varied. The uplift rate in the Nanga-Parbat-Haramosh
massif, for example, increased from :;0.5 to more than 10
451
mm1a during the past 7 Ma (Hurford 1991; Burbank et al.
1996). In the backthrust zone of the Transhimalayas with
exposed sedimentary rocks, the long-term (20 Ma) average
rate of denudation was 100-200 mmlka only (cf. Sect.
11.5.5).
Modern rates of tectonic uplift are sometimes unusually high because of delayed isostatic rebound in response to ice melted after the last glacial period.
Evaluating repeated geodetic surveys or dated elevated fluvial terraces, very high uplift rates (~3
mm1a) were found in several high mountain ranges
(Gansser 1984). Using geomorphological marker horizons, Darmendrail et al. (1993) found uplift rates of
1 to 3 mrnIa for the past 100 ka in the French Alps.
In the less glaciated Umbrian Apennines, central
Italy, uplift rates in the Quatemary ranged from ~250
to 500 mmlka (Rasse 1993). More data ofuplift rates
in various tectonic settings are summarized in special
articles (e.g. Leeder 1991).
9.6.3 Simple Uplift-Denudation Models
Substantial denudation is commonly initiated by tectonic and· thermal uplift (both included in the term
"tectonic" uplift in the following). Later, after the
end of this process, rock uplift is maintained to some
extent by isostatic rebound following erosional unloading as mentioned above (Fig. 9. 17a). Repeated or
changing vertical tectonic motions and their effects
on denudation and landscape evolution are demonstrated in the models shown in Fig. 9.18a and b and
Fig.9.151:
- Scenario 1, teetonie pulses (Fig. 9.18a). Lowland
(around base level) is raised rapidly by tectonic uplift
during a short time period. After this event, landscape lowering is controlled by denudation and isostatic rebound until a second tectonic pulse increases
mean elevation and thus generates a second period of
slow, exponential decay of the land surface. In this
case, no equilibrium between the (changing) denudation rate and total uplift during the pulses is reached
because the response time of the system is longer
than the tectonic pulse ("Penck-type" landscape evolution; see e.g. Kooi and Beaumont 1996). Finally, a
short period of subsidence brings about accelerated
surface lowering followed again by slow decay.
- Scenario 2, eontinuous lang-term upliji (Fig.
9.18b). Rapidly changing rates of epeirogenic uplift
force denudation to follow until, after some time, a
kind of dynamic equilibrium (steady-state condition)
between total uplift and denudation is obtained
("Hack-type" landscapes). After each change in the
rate of uplift, the system approaches a new dynamic
equilibrium at higher or lower elevation above base
level. The equilibrium level also follows a lowered
base level with a certain response time, t c •