Available online at www.sciencedirect.com
Lithos 102 (2008) 1 – 11
www.elsevier.com/locate/lithos
The influence of lithospheric thickness variations
on continental evolution
Dan M c Kenzie ⁎, Keith Priestley
Institute of Theoretical Geophysics, Bullard Labs of the Department of Earth Sciences, Cambridge University, UK
Received 21 November 2006; accepted 22 May 2007
Available online 6 June 2007
Abstract
The shear wave velocity Vs as a function of depth z can be obtained from surface wave tomography, using the phase velocities
of fundamental and higher mode Rayleigh waves. Since Vs is principally controlled by temperature, rather than by composition, it
can be used to map the lithospheric thickness. Extensive regions of thick lithosphere underlie some, but not all, cratons.
Conversely, thick lithosphere underlies some platforms and belts of active deformation. Because of this lack of correspondence,
and because their age cannot be determined from seismology, we refer to regions of thick lithosphere as ‘cores’ rather than
‘cratons’. The shape of such cores has controlled the geometry of continental deformation and the distribution of diamond-bearing
kimberlites. The strength of the cores resides in the dry crust, which is insulated from the hot convecting mantle by the thick
buoyant lithosphere. The most surprising feature is the presence of thick lithosphere beneath Tibet and Iran, whose velocity
structure closely resembles that of the cores beneath cratons, though they have a thicker hotter crust. Tibet and Iran appear to be
places where cratons are now being formed.
© 2007 Elsevier B.V. All rights reserved.
Keywords: Surface wave tomography; Continental geotherms; Tibet; Iran; Formation of cratons; Lithosphere
1. Introduction
For at least a hundred years geologists concerned with
continental tectonics have recognised the importance of
the rigidity of shields and cratons (see reviews by Suess
(1909) and Holmes (1965)). Once plate tectonics had been
generally accepted, it became clear that the properties of
the entire lithosphere beneath cratons were likely to be
responsible for their behaviour, rather than those of the
crust alone. Detailed studies of the mineralogy of mantle
nodules, especially those from diamond bearing kimberlites, have shown that the lithospheric thickness beneath
⁎ Corresponding author. Tel.: +44 1223337191.
E-mail address: [email protected] (D. McKenzie).
0024-4937/$ - see front matter © 2007 Elsevier B.V. All rights reserved.
doi:10.1016/j.lithos.2007.05.005
some cratons is as great as 250 km, (see Finnerty and
Boyd, 1987). However, until recently, there was no
method of mapping the three dimensional structure of
cratonic lithosphere, partly because large areas of cratons
are covered with thin veneers of Phanerozoic sediments. It
is then difficult to use geological observations to map the
margins of cratons.
2. Determination of lithospheric thickness
As so often happens in the Earth Sciences, an advance
in technology has produced the key to an old question. In
this case the key has been provided by digital
seismometers, which have provided broadband records
of the ground motion, especially of surface waves.
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D. M cKenzie, K. Priestley / Lithos 102 (2008) 1–11
The velocity structure of the upper mantle affects the
phase velocity of fundamental and higher mode surface
waves, and also the travel times of P and S waves from
teleseismic sources. Priestley and McKenzie (2002)
discuss in detail what constraints such measurements
provide on the horizontal and vertical variations in
seismic velocity. Because teleseismic ray paths are steeply
inclined, the resulting models have excellent lateral
resolution but poor vertical resolution. In particular,
most of the methods used to invert such measurements
remove the average vertical velocity structure, which is
therefore unconstrained (see Priestley and McKenzie,
2002). In contrast, surface wave tomography has limited
lateral resolution, of 300–400 km in the models discussed
below, but excellent vertical resolution, of 30–50 km,
especially if both fundamental and higher mode phase
velocities are used in the inversion (Debayle and Kennett,
2000; Ritsema and van Heijst, 2000; Priestley and
Debayle, 2003). Because of these differences, velocity
models based on teleseismic travel time tomography do
not in general agree with those obtained from surface
wave tomography. In particular, the regional lithospheric
thickness cannot be estimated from teleseismic travel time
tomography if the influence of the horizontal average of
velocity structure is removed during the processing.
From a geological point of view, maps of the variation
of lithospheric thickness are more useful than those of Vs
(z). But it is not straightforward to estimate the
lithospheric thickness from Vs(z) or Vp(z). The problem
is that the base of the lithosphere does not correspond to a
step change in temperature or composition, and is therefore not associated with a change in velocity. Therefore in
general no conversion occurs, from P to S or from S to P,
when seismic waves cross the base of the lithosphere,
which cannot be mapped using receiver functions. What
does change rapidly over a limited vertical distance at the
base of the lithosphere is the temperature gradient. Since
Vs is more sensitive to temperature than is Vp, and since
the phase velocities of surface waves are more sensitive to
the S wave velocity than they are to that of P, Priestley and
McKenzie (2006) used Vs(z) to estimate the lithospheric
thickness. They first parameterised the dependence of Vs
on pressure and temperature using
Vs ¼ Vs ð P; h; aÞ
ð1Þ
where P is the pressure, θ the temperature in °C and a is a
variable describing the activation process
a ¼ A Vexp ð E þ PVa Þ=RT
ð2Þ
where A′is a frequency factor, E the activation energy,
Va the activation volume and T is the temperature in
Kelvin. It is convenient to remove the nonactivated part
of the pressure dependence of Vs by writing
Vs4 ¼ Vs =ð1 þ bv ð z 50ÞÞ
ð3Þ
where z is the depth in km and bV is an empirical
constant. Then
Vs4 ¼ mh þ c þ A exp ½ð E þ PVa Þ=RT ð4Þ
The values of the constants obtained by Priestley and
McKenzie are
bV ¼ 3:84 104 km1 ;
m ¼ 2:8 104 km s1 -C1
c ¼ 4:72km s1 ;
A ¼ 1:8 1013 km s1
E ¼ 409kJ mol1 ;
Va ¼ 10 106 m3 mol1 u10cm3 mol1
ð5Þ
Eqs. (3) and (4) can be inverted to give T (Vs) using
Newton-Raphson iteration. In continental regions, values
of Vs(z) were used to obtain T(z) at intervals of 2° in
latitude and longitude, and at depth intervals of 25 km.
These profiles were then fitted to a geotherm, calculated
using the methods outlined by McKenzie et al. (2005), and
using an interior potential temperature of 1315 °C. The
average potential temperature of the mantle is constrained
by the average thickness of the oceanic crust of 7 km,
whose thickness changes by 1 km when the potential
temperature changes by 13 °C (McKenzie et al., 2005).
Because the oceanic crustal thickness is independent of
location and of the age, and only has a standard deviation
of about 1 km, the potential temperature of the mantle
is constrained to be 1315 ± 13 °C. McKenzie et al.'s
approach assumes that the thermal structure of the
lithosphere is controlled by local processes; heat generation in the crust, transport of heat by conduction through
the mechanical boundary layer and by advection and
conduction in the thermal boundary layer below. They
assumed that it is the large scale mantle circulation that
keeps the average potential temperature beneath oceans
and continents constant. Recent large scale modeling of
mantle convection in the presence of thick undeformable
continental lithosphere (Cooper et al., 2004) generates
continental geotherms that agree with those calculated by
McKenzie et al. (2005) if the thermal conductivity is taken
to be independent of temperature and the mantle is
assumed to be incompressible, when the potential and real
temperatures are the same. Both McKenzie et al. (2005)
and Cooper et al. (2004) assume that the depleted shallow
D. M cKenzie, K. Priestley / Lithos 102 (2008) 1–11
part of the continental lithosphere remains attached to the
continental crust and is not removed by the small scale
convective instabilities that control the thickness of the
thermal boundary layer. Whether this behaviour occurs
depends on the difference in density between depleted and
fertile peridotite and on the viscosity of the thermal boundary layer. Lenardic and Moresi (1999), Sleep (2003) and
Lee et al. (2005) have investigated this question, and
argue that the depleted layer is likely to remain attached.
For various reasons estimates of T(z) obtained from
Vs(z) are more accurate at high temperatures near the
base of the lithosphere than they are near the Moho. The
velocities at depths of less than 100 km are lower than
their true values because they are influenced by the lower
velocities of the crust. This effect causes the estimated
temperatures to be higher than the true values. The
temperature estimates are more accurate at high
temperatures than they are at low temperatures, because
|∂Vs/∂T| increases with increasing temperature. Therefore only those values of T(z) that exceeded 1100 °C at
depths of 125 km or more were used to obtain the
geotherms and to estimate the lithospheric thickness.
Priestley and McKenzie (2006) also calculated profiles
of Vs(z) for their thermal models, most of which show no
obvious change in Vs at the base of the lithosphere where
there is a change in the temperature gradient. The
geotherms obtained from Vs(z) were used to estimate the
depth of the graphite–diamond phase change and the
thickness of the lithosphere in the diamond stability
field, using the phase boundary given by Kennedy and
Kennedy (1976). The depth of the graphite–diamond
phase change is about 140 km where the lithospheric
thickness exceeds 200 km. Estimates of lithospheric
thickness from Vs and from the nodule mineralogy (Brey
and Kohler, 1990) are independent and agree to within
about 25 km (Priestley and McKenzie, 2006). Figs. 1–5
show contours of the lithospheric thickness beneath all
continents except South America and Antarctica. The
path coverage of these two continents is not yet sufficient
to allow satisfactory maps to be produced.
3. Lithospheric thickness variations
If diamonds in kimberlites are lithospheric xenocrysts,
they should only be present where kimberlites erupt
through regions where the lower part of the lithosphere is
in the diamond stability field. The map of lithospheric
thickness beneath Africa (Fig. 1) shows that most of the
diamond-bearing kimberlites satisfy this condition. Fig. 1
suggests that diamondiferous kimberlites tend to occur on
the margins of the regions of thick lithosphere, rather than
in the centres. Such a distribution is to be expected,
3
Fig. 1. Contours of the thickness of the African lithosphere, calculated
from the shear wave velocity. The small back dots show the locations
at which this thickness is calculated. The magenta circles show the
locations of diamond-bearing kimberlites (Nixon, 1987; Janse and
Sheahan, 1995), and the yellow circles show those of alkali basalts
containing mantle nodules whose mineral compositions have been
used to estimate the lithospheric thickness. The numbers in white
boxes show the thickness of the lithosphere estimated from
the composition of minerals in mantle nodules (see Priestley and
McKenzie (2006) for references). The cratons are outlined in yellow
and are labelled as follows: a West African Craton, b Angolan Craton,
c Tanzanian Craton, d Kalahari Craton.
because diamond-bearing kimberlites require both melt
and diamond-bearing lithosphere. Since the solidus
temperature increases with pressure, less melt is likely
to be generated beneath the thickest lithosphere than
beneath the thinner regions on either side.
The geometry of the thick parts of the lithosphere in
Fig. 1 is somewhat different from that proposed on geological grounds, especially in southern Africa (see Nixon,
1987, Fig. 98). Fig. 1 shows that the Kalahari, Tanzanian
and Angolan Cratons are not now separate structures, but
all form part of a larger region of thick lithosphere. Such
continuous regions of thick lithosphere are hereafter called
‘cores’. The two in Fig. 1 will be referred to as the West
African and Southern African Cores.
Different parts of what is now a single core may have
had very different histories, and may only have been
joined together relatively recently. The value of Vs
provides no age information, so there is no a priori reason
to believe that all the rocks that form a core were formed at
the same time as the cratons that many of them contain.
4. Continental tectonics
Maps of the continental cores in Figs. 1–5 show that
they have controlled the geometry of the continental fold
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D. M cKenzie, K. Priestley / Lithos 102 (2008) 1–11
Fig. 2. (a) Tectonic map of North America (Holmes, 1965, Fig. 811).
(b) As for Fig. 1. The two green lines marked COC and SC show the
locations of the COCORP profile and the southern Canadian sections
discussed in the text. NM marks the location of the New Madrid
earthquake sequence. The yellow line shows the approximate
boundary of the North American Craton, labelled a.
belts. The North American Core (Fig. 2) is fringed by the
Appalachians in the east and the Rocky Mountains in the
west. In the southeast, a COCORP seismic profile (Cook
et al., 1979) shows that a thin sheet of younger rocks
have been thrust over the Core. Similar behaviour occurs
in the west, near 50°N in southern Canada, where the
southern Rocky Mountains override the Core in the
classic area of thin skinned tectonics (Bally et al., 1966).
As in Africa, most of the diamond-bearing kimberlites
occur within the North American Core. Fig. 2 also shows
that the enigmatic New Madrid earthquakes occurred at
the margin of the Core.
The western part of the North Eurasian Core (Fig. 3)
has controlled the tectonic evolution of Europe, where the
Caledonian and Hercynian fold belts wrap round its
margins. In Norway the Caledonian thrusts override the
margins of the Core, as do the Hercynian and younger
thrusts in southern Russia and the Caucasus. The Core
consists of the Baltic Shield and the Russian and Siberian
Platforms, and the part west of the Urals is one of the
thickest cores illustrated in Figs. 1–5. Three small regions
of thick lithosphere are also present in Fig. 3. One of these,
south of the Aegean, is unlikely to be real. It is probably
caused by the high velocity in the slab that is being
subducted by the Hellenic arc. The other two, beneath the
British Isles and Iran, are likely to be real. The Urals
themselves cut across the Core, which therefore must
have been assembled after the Carboniferous.
The south-east boundary of the North Eurasian Core
(Fig. 4a) controls the northern edge of the deformation
associated with the Himalaya and the northward motion
of India. This control is especially obvious east of Lake
Baykal. It is surprising that the lithosphere is so thin
throughout most of China and beneath the southern part
of the Indian Peninsula. Eastern Chinese diamondbearing kimberlites of Ordovician age occur in a region
where the lithosphere is now thin (Menzies et al., 1993;
Griffin et al., 1998). Their presence suggests that the
lithosphere has been thinned since the Ordovician. The
Indian kimberlites are Precambrian, and the southern
part of the Indian shield may also have been thinned
since they were emplaced.
The maps in Figs. 2–5 also provide information about
how large cores form by joining smaller ones together,
and how single cores can divide into two or more pieces.
The Urals cross the North Eurasian Core, yet the region
of thick lithosphere is continuous between the Russian
and Siberian Platforms. Since it is unlikely that the edges
of these platforms matched before they came together,
the edges of cores must be deformable, or the thinner
lithosphere that existed between the cores must have
been thickened by shortening. Fig. 4b shows a reconstruction of the North Atlantic, using the pole that
Bullard et al. (1965) obtained by fitting the 500 fathom
contours of Africa and North America. The edges of the
North American and West African Cores are parallel, and
probably formed a single core before the two continents
separated. The thin lithosphere that is now present
on both margins is likely to have been produced by
stretching before seafloor production began, and is likely
to consist in part of depleted harzburgite. The break
follows the Caledonian belt, which formed a line of
D. M cKenzie, K. Priestley / Lithos 102 (2008) 1–11
5
Fig. 3. (a) Tectonic map of Europe (Holmes, 1965, Fig. 807). (b) As for Fig. 1. The yellow line shows the approximate boundary of the East European
Craton, labelled a.
weakness. Even earlier, Caledonian shortening closed
the Iapetus Ocean. Like the Urals, this closure only
stopped when the thin lithosphere between the cores had
been removed by thickening or lateral transport. Fig. 5
shows that the Australian Core also has a considerably
greater extent than the exposed shields.
5. Underlying physical processes
Nodules from kimberlites show that much of the
mantle part of cores consists of depleted harzburgite,
with a density that is about 40 kg/m less than that of the
fertile mantle. The seismic velocity change associated
with this depletion is only about 0.03 km/s (Jordan,
1979). Therefore the density of the depleted material is
controlled principally by its composition, and its velocity
by its temperature (Priestley and McKenzie, 2006). The
low density of the harzburgite stabilises the cores against
convective instabilities, which reduces their steady state
temperature gradients. Maggi et al. (2000) argued that
the strength of old continental lithosphere results from
the strength of the crust, not the mantle. McKenzie et al.
(2005) showed that the distribution of earthquakes with
depth and the variations in the thickness of the elastic
layer that supports long term lithospheric stresses both
require the elastic stresses to be relaxed by creep when
the mantle temperature exceeds about 600 °C. This result
requires the strength of the continental lithosphere to
reside in the crust, not the mantle. Furthermore, if the
crust is to be strong when its temperature is as high as
600 °C, it must be dry (Mackwell et al., 1998; Jackson
et al., 2004). The striking influence of the cores on
continental deformation must then be a consequence of
the thermal insulation of the crust by the buoyancy of the
thick stagnant layer of mantle below. Resistance to
lithospheric shortening will therefore increase sharply
6
D. M cKenzie, K. Priestley / Lithos 102 (2008) 1–11
Fig. 4. (a) As for Fig. 1. The cratons are outlined in yellow and are labelled as follows: a East European Craton, b Siberian Craton, c North China
Craton, d Yangtze Block of the South China Craton, e Bastar, Singhbhum, and Arivalli Cratons, and f Dharwar Craton. (b) Reconstruction of the
Central Atlantic using the fitting pole of Bullard et al. (1965).
when approaching cores meet, largely because of the
stresses required to deform crust which is dry and relatively cold. Conversely, stretching of a core, especially
when localised by an old line of weakness, can rapidly
thin the lithosphere, and hence decrease its strength.
Because the shear wave velocity of mantle material is
principally controlled by temperature rather than by
composition, the seismic velocities of stretched cores,
consisting of depleted harzburgite, will be the same as
those of fertile mantle. So the buoyant depleted material
can only be distinguished from fertile mantle where it
forms thick cold roots, and there is as yet no method of
knowing what parts of the mantle part of lithosphere of
normal thickness (∼ 100 km) consist of depleted mantle.
The thin lithosphere beneath the diamond-bearing
kimberlites of NE China (Fig. 4a) may still contain
harzburgite, and perhaps the same is true of southern
India. As well as controlling the tectonic deformation of
continents, the geometry of cores is also likely to control
the composition and abundance of alkaline magmas
D. M cKenzie, K. Priestley / Lithos 102 (2008) 1–11
7
Fig. 5. As for Fig. 1. The cratons are outlined in yellow and are labelled as follows: a Darwin Craton, b Kimberley Craton, c Pilbara Block, d Yilgarn
Block, e Gawler Craton.
generated near the base of the lithosphere. The parameters
that are most sensitive to the depth of melt generation are
the concentrations of sodium and potassium. Because the
compatibility of sodium in clinopyroxene changes with
increasing pressure, from being incompatible at depths of
less then about 70 km to being compatible at 150 km
(Wang and Takahashi, 1999; Tsuruta and Takahashi,
1998), alkaline magmas become increasingly potassic as
the depth of melting increases. Measurements of alkali
concentrations are routinely reported, and therefore
provide a simple method of testing this suggestion.
6. Craton formation
The thickness of 75 Ma old oceanic lithosphere is
about 100 km. Thereafter its thickness does not increase
with increasing age. Though a number of explanations
for this behaviour have been proposed, it is now widely
believed that it results from a convective instability that
removes the lower part of the lithosphere. Since the
composition of the convecting mantle beneath the plates
cannot be affected by crustal thickness variations, the
same convective instability should occur beneath
continents, especially if the entire lithosphere has been
thickened by thrusting. This argument led Houseman
et al. (1981) to propose that the shortening that increased
the crustal thickness beneath Tibet should produce an
unstable lithosphere whose lower part would detach and
sink into the mantle. They argued that this behaviour
was required to account for the regional metamorphism
which has occurred beneath Tibet since India and Asia
collided in the Tertiary. Their proposal has been used to
account for a variety of geophysical and geological
observations, such as the uplift of Tibet and the low
values of Vs observed in the mantle close to the Moho
beneath northern Tibet (McNamara et al., 1995).
However, the regional maps of Vs(z) and of lithospheric thickness beneath Tibet and Iran (Fig. 4a) are not
compatible with Houseman et al.'s proposal. They show
that the thickness of the lithosphere beneath both plateaus
is now about 260 km. This result is a surprise, and
suggests that the process that generated the thickened
crust beneath these plateaus has also resulted in thick
lithosphere that extends beneath the whole plateau. The
lateral resolution of the surface wave tomography is about
400 km, so narrow vertical high velocity features like that
described by Tilmann et al. (2003) will not be resolved.
The centre of the plateau has been elevated for the last
35 Ma (Rowley and Currie, 2006) and is still underlain by
thick lithosphere. There is therefore no evidence of large
scale lithospheric removal at any time since the formation
of the plateau. The thick lithosphere must be stabilized
against convective instability, most probably by depletion. Melt removal from fertile upper mantle produces a
low density residue, but scarcely affects the value of Vs.
The resemblance of the lithospheric structure of
Tibet to that of Archaean and Proterozoic cratons is
unexpected. The obvious explanation is that both are
produced by the same processes, and that Tibet may
eventually become a craton when the upper 30–40 km
of its crust has been removed by erosion. The crustal
thickness of most cratons is now about 40 km, and the
8
D. M cKenzie, K. Priestley / Lithos 102 (2008) 1–11
Fig. 6. Sketch showing two processes that can produce thick crust
overlying thick lithosphere, by (a) uniform thickening of crust and
lithosphere, initially both of normal thickness, and (b) thickening the
crust by emplacing a thick crustal layer on top of a craton, whose crust
is initially of normal thickness, but whose lithospheric thickness is
200–250 km.
metamorphic rocks at their surface record pressures of
about 1 GPa (Harley, 1989, 2004). Their crustal
thickness must therefore have been 70–80 km at the
time the metamorphism occurred if they have not since
been deformed. Presumably such large crustal thicknesses resulted from shortening. An important question
is whether the whole lithosphere was also shortened, or
whether only the crust was thickened, by thrusting or
lower crustal flow. Seismic studies (Tilmann et al.,
2003) suggest that the crust and lithosphere of the Indian
Shield extends beneath the southern part of Tibet, and
that the thick lithosphere in this region is inherited. This
region therefore does not provide information about
how thick lithosphere beneath shields is produced.
However in northern Tibet Haines et al. (2003) have
argued that the Tibetan crust has been shortened by pure
shear. This process could thicken the lithosphere if the
mantle was deformed in the same way.
Fig. 7. Evolution of the lithospheric temperature (a) and shear wave velocity (b), calculated from Eqs. (3) and (4), after sudden thickening of the
lithosphere. Before shortening the lithospheric thickness was 80 km, and the upper and lower crustal thickness were 20 km and 5 km, with heat
generation rates of 2 μW/m3 and 0.4 μW/m3 respectively. The lithosphere was uniformly shortened by a factor of 3 at t = 0. The thermal conductivity
of the crust has the same temperature dependence as that of the mantle, but its value is half that of the mantle. The green line in (a) shows the location
of the Moho. The velocities in (b) are labelled with the value of t in Ma, and are incorrect in the crust.
D. M cKenzie, K. Priestley / Lithos 102 (2008) 1–11
Many of the nodules extracted from thick lithosphere
by kimberlites are harzburgites that have been produced
from fertile mantle by the extraction of about 30% melt.
Canil and Wei (1992) and Canil (2004) have argued that
the Cr concentration in these rocks requires the
depletion to have occurred in the presence of spinel,
because it is the only commonly occurring phase in
peridotites in which Cr is compatible. This argument
requires the melt to have been removed at pressures of
less than about 3 GPa, or at depths of less than 100 km
(Canil, 2004). Since harzburgite nodules come from
depths as great as 200 km, the thick lithosphere must
have been produced by shortening after melt extraction.
Tainton and McKenzie (1994) argued that the extent of
depletion also requires thick lithosphere to have been
produced by shortening, because no choice of mantle
potential temperature can generate 30% melt at depths
as great as 150 km. Both arguments therefore require
thick depleted lithosphere to have been produced by
shortening, to transport the depleted material downwards to form the cratonic roots.
If lithospheric removal has not occurred beneath
Tibet, the heat responsible for regional metamorphism
cannot have been conducted in from the base of the crust,
and a different process must also be responsible for the
low sub-Moho velocities. The only other heat source
available is crustal radioactivity, and it is clearly of
interest to discover whether it can produce the observed
effects. Fig. 6 shows a sketch of two processes that can
generate thick crust and lithosphere like that beneath
Tibet; by shortening the entire lithosphere, or by increasing the crustal thickness of a region whose lithospheric thickness is already 200 km or more. Fig. 7 shows
the thermal evolution of lithosphere formed by the first
of these processes, starting with a lithospheric thickness
of 80 km and a crustal thickness of 25 km. This lithosphere is instantaneously shortened by a factor of
three at t = 0. As England and Thompson (1984) and Le
Pichon et al. (1997) have pointed out, crustal radioactivity is itself sufficient to cause the temperature at midcrustal levels to rise rapidly. Fig. 7a shows its temperature reaches 900 °C 60 Ma after the shortening
event. Fig. 7b shows that conduction of heat downwards
from the crust causes Vs of the uppermost mantle beneath
the Moho to decrease. Comparison of the behaviour of
the two models in Fig. 6 shows that radioactive heat
generation in the thickened crust, and not the process by
which the thickening occurs, is the principal control on
the thermal behaviour. Since the crustal temperatures
exceed the granite solidus, melting will occur, and the
upward movement of melt will transport the elements
that generate heat to the upper part of the crust.
9
The final stage of this process occurs after shortening
ceases, when erosion removes the upper crust and its
radioactive elements, leaving the depleted mid-crust that
underwent granulite facies metamorphism exposed at
sea level. This process does not change the thickness of
the mantle part of the lithosphere, and does not involve
any tectonic deformation. Fig. 8a shows a number of
estimates of pressure and temperature from granulite
Fig. 8. (a) Temperature and pressure from the model in Fig. 6, labelled
with the value of t in Ma. The pressure and temperature estimates for
granulites in the inset box are from Harley (1989). The red box shows
similar estimates from Tibetan crustal xenoliths from Hacker et al.
(2000). (b) as for (a) with the same concentrations of K, U and Th, but
with the shortening taking place 3 Ga ago.
10
D. M cKenzie, K. Priestley / Lithos 102 (2008) 1–11
terranes from Harley (1989), together with those from
Tibetan crustal xenoliths (Hacker et al., 2000). The plot
shows that the geothermal evolution in Fig. 7a can
produce mid-crustal temperatures of 900–1000 °C
required by the observations. In the past the rate of
radioactive heating was greater. Fig. 8b shows the
thermal behaviour corresponding to that in Figs. 7 and 8a
if the shortening occurred 3 Ga ago, calculated using the
present concentrations of K, U and Th. As expected, the
temperatures are higher than those in Fig. 8a, and exceed
the highest temperatures estimated from the mineralogy
of ultrahigh temperature metamorphic terranes (Harley,
2004).
These numerical experiments show that crustal
radioactivity can produce granulite grade regional metamorphism and lower the value of Vs in the mantle beneath
the Moho. They also show how processes that are now
occurring can account for the main features of cratons.
7. Conclusions
Regions of thick, cold continental lithosphere can
now be mapped using surface wave tomography. The
resulting maps show that a number of individual, mostly
Precambrian, shields are now parts of larger regions of
continuous thick lithosphere. We propose that such
regions are called ‘cores’, partly to distinguish them from
the individual shields and cratons that have been joined
by tectonic processes to form the cores, partly because
their age of formation cannot be determined from
seismology, and partly because not all cratons are now
underlain by thick lithosphere. These cores have
dominated the tectonic deformation of the continents,
and control the distribution of diamond bearing
kimberlites. Their margins can be deformed, and are
often overridden by younger thrusts. Cores are difficult
to shorten, but are easier to stretch, especially on preexisting lines of weakness. Their strength largely resides
in their dry crust, which remains relatively cold because
it is insulated from the convecting mantle by a thick layer
of low density harzburgite. The Tibetan Core now has a
thickness of 250–300 km, and has not been thinned by
delamination. If the upper half of the Tibetan crust is
removed by erosion, the resulting velocity structure will
be similar to that of cores whose surface is now at sea
level, many of which were formed in the Archaean.
Acknowledgements
We would like to thank C. Lee, J. Jackson, R. Rudnick
and N. Sleep for their help, and the Royal Society for
support.
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