crustal collapse, mantle upwelling, and cenozoic · extension in the cordillera, including...

TECTONICS, VOL. 17, NO. 2, PAGES 311-321, APRIL 1998

Crustal collapse, mantle upwelling, and Cenozoic extension in the North American Cordillera

Mian Liu and Yunqing Shen Department of Geological Sciences, University of Missouri, Columbia

Abstract. Gravitational collapse has been suggested as the major cause of Cenozoic extension in the North American Cordillera and many other orogenic belts. Although both crustal thickening and mantle upwelling may have contributed to the Cordilleran extension, previous models of gravitational collapse have focused on the former; the cause of mantle upwelling and its relationship to crustal collapse remain obscure. Here we attempt to address the question of whether gravitational collapse of an overthickened crust could induce major mantle upwelling and whole-lithosphere extension. Thermal-rheological calculations indicate that crustal collapse may decouple from the mantle lithosphere, because the extensional forces arising from an overthickened crust are limited to the crust, while the rheology of continental lithosphere is intrinsically stratified. Even when the mantle lithosphere is mechanically coupled to the crust, thermomechanical modeling indicates that strain is localized in the weak lower crust during crustal collapse, and no significant (<10 km) thinning of the mantle lithosphere may be induced at the absence of extensional forces from plate boundaries. Crustal collapse of the Sevier-Laramide orogen seems adequate to account for much of the mid-Tertiary extension in the Cordillera, including formation of many core complexes, but it is unlikely to have been the major cause of the more recent basin-and-range extension. We suggest that a strong pulse of mantle upwelling in the mid-Tertiary, as indicated by the "ignimbrite flare-up," may have triggered basin-and-range extension by weakening the lithosphere and providing excess gravitational potential energy. The cause of mantle upwelling remains uncertain, but the continued extension and volcanism since mid-Miocene in the northern

Basin and Range province favor an active mantle upwelling with internal convective heating.

1. Introduction

Orogenic belts, the products of intensive compressional tectonics, are often the sites of continental extension and

rifting, as first observed by Wilson [1966] in his study of the opening of the Atlantic Ocean. The theory of plate tectonics offers no ready explanation for such extension, because orogenic belts usually form near convergent plate boundaries, and extension of orogenic belts happens even when plates are

Copyright 1998 by the American Geophysical Union.

Paper number 98TC00313. 0278-7407/98/98TC-00313512.00

311

still converging and the regional stress field is predominantly compressional. This is illustrated by active extension in the central Andes [Dalmayer and Molnar, 1981; Dewey, 1988] and the Tibetan plateau [England and Houseman, 1989; Molnar and Chen, 1983].

The popular explanation is the hypothesis of gravitational collapse [Dewey, 1988; Molnar and Lyon-Caen, 1988]. Most orogenic belts in the world are supported by an Airy-type crustal root [Airy, 1855]. The isostatic balance of vertical forces, however, does not mean mechanical equilibrium in the lateral directions. A thickened crust, with a higher gravitational potential energy than the adjacent lowlands, tends to spread or collapse under its own weight [Artyushkov, 1973; McKenzie, 1972]. Gravitational collapse is now believed to have played a major role in both synorogenic and postorogenic extension in many orogenic belts [Dewey, 1988], including the Cenozoic extension in the North American Cordillera (Figure 1) [Coney, 1987; Harry et al., 1993; Livaccari, 1991; Sonder et al., 1987; Wernicke et al., 1987].

Although gravitational collapse of an overthickened crust is conceptually simple, its geodynamics are not well understood. One major problem is the role of mantle upwelling, which is often involved in orogenic collapse and contributes to the driving forces [Dewey, 1988; England and Houseman, 1989]. However, because mantle processes are difficult to constrain, most studies have avoided them or

included them in calculations without offering much discussion of their cause and relationship to crustal collapse [Sonder et al., 1987]. Various assumptions have been made, leading to divergent conclusions. Some workers, assuming mechanical decoupling of the crustal processes from the mantle lithosphere because of the weak lower crust, have predicted quick flattening of crustal welts into "pancakes" with little effect on the mantle lithosphere [Bird, 1991]; others, assuming a full crust-mantle coupling, have shown that extensional collapse of an overthickened crust can lead to major extension of the whole lithosphere [Govers and Wortel, 1993; Harry et al., 1993].

In this study we attempt to constrain the role of crustal collapse and the relationship between crustal collapse and mantle upwelling. The well-established history of the Cenozoic extension in the North American Cordillera

provides a good example for this study. The questions we hope to address include the following: (1) What is the role of gravitational collapse of the Sevier-Laramide orogen in the Cordilleran extension? (2) Could crustal collapse of the Sevier-Laramide orogen have led to basin-and-range extension? (3) What is the cause of mantle upwelling under the Basin and Range province?

312 LIU AND SHEN: CRUSTAL COLLAPSE AND CORDILLERAN EXTENSION

50 ø

40 ø

30 ø

125 ø 115 ø 105 ø

Figure 1. Shaded relief map of the North American Cordillera. The tooth curve shows the eastern boundary of the Sevier-Laramide fold-and-thrust belts. The bold solid line indicates the boundary of the Basin and Range province. The stippled areas are the locations of major metamorphic core complexes, and the thin solid line shows the main volcanic field of the ignimbrite flare-up during mid-Tertiary. Abbreviations are as follows: NBR, northern Basin and Range, also called the Great Basin (GB); SRB, southern Basin and Range; SRP, Snake River Plain; and CP, Colorado Plateau.

2. Cenozoic Extension in the North American Cordillera

There are numerous excellent review articles on extensional

tectonics in the North American Cordillera [Burchfiel et al., 1992; Burchfiel and Davis, 1975; Coney, 1987' Eaton, 1982; Hamilton and Myers, 1966; Stewart, 1978' Thompson and Burke, 1974; Wernicke, 1992; Zoback et al., 1981]. A summary here highlights some of the problems pertinent to our discussion.

During the Mesozoic and early Cenozoic (-165-55 Ma), western No_rth America experienced a protracted phase of crustal compression as the oceanic Farallon plate subducted underneath North America. The crustal contraction involved a

complex history of subduction-related deformation and massive plutonism along the coastal margin [Burchfiel and Davis, 1975]. Crustal shortening in the inland Cordillera occurred in a zone stretching from Canada to northern Mexico; it telescoped more than 200 km of crust and caused progressive development of the fold-and-thrust belt [Elison, 1991]. In the hinterland of the orogenic belt the crustal thickness was nearly doubled to more than 50 km [Coney and Harms, 1984; Parrish et al., 1988].

Although localized extension may have started as early as the Mesozoic [Hodges and Walker, 1992], major extension in the Cordillera is postorogenic, occurring after the Laramide orogeny [Zoback et al., 1981]. In many places the inception of extension is marked by development of metamorphic core complexes [Coney, 1987]. The occurrence of many core complexes along the core zone of the Sevier-Laramide orogen where crust was significantly thickened strongly indicates a cause of crustal collapse [Coney and Harms, 1984]; however, not all core complexes are formed in the overthickened core zone. In the southern Basin and Range province, most core complexes developed in the midst of a deep-seated thrust belt [Coney, 1980]; detachment faults and core complexes are also found in the Mojave desert where there is no clear evidence for a significantly thickened crust before extension [Dokka and Ross, 1995; Glazner and Bartley, 1984]. Furthermore, the inception of major extension in the Cordillera was diachronous, although the duration and the amount of crustal contraction were remarkably uniform from southeastern British Columbia to Nevada and Utah [Elison, 1991]. In the southern Canadian Cordillera, northern Washington, Idaho, and Montana, extension began in early Eocene. Farther south, the inception time was largely Oligocene in the Great Basin

LIU AND SHEN: CRUSTAL COLLAPSE AND CORDILLERAN EXTENSION 313

(the northern Basin and Range province), and was slightly later in the Mojave-Sonora desert region. Extension in areas near the latitude of Las Vegas did not occur until mid-Miocene [Wernicke et al., 1987]. Wernicke et al. [1987] have pointed out that the timing of inception of extension in the Cordillera is apparently correlated with the abundance of associated plutonism, and numerous studies have stressed the importance of magmatism in core-complex formation [Armstrong and Ward, 1991;Axen et al., 1993]. Much of the magmatism may have resulted from postorogenic thermal relaxation and radioactive heating [Glazner and Bartley, 1985]; however, at least in the Canadian Cordillera where extension occurred

within a few million years after the orogeny, some thermal perturbations from the mantle seem necessary [Liu and Furlong, 1993].

Calc-alkaline volcanism was widespread through much of the early to middle Tertiary and culminated with eruption of voluminous (>35,000 km 3) silicic tuff in the central Great Basin between 34 and 17 Ma [Best and Christiansen, 1991]. Although eruption of mafic magma is rare, significant mantle upwelling seems necessary to provide heat for the extensive crustal anatexis [Hildreth, 1981] and source materials for some of the silicic tuff [Grunder, 1995; Johnson, 1991]. The relationship between extensional tectonics and volcanism is controversial. Some workers find that extension was mainly synvolcanical [Gans et al., 1989], others [Axen et at., 1993; Best and Christiansen, 1991; Taylor and Bartley, 1992] argue for a poor spatial-temporal correlation between volcanism and extension on a provincial scale. Liu and Furlong [1994] suggested that the apparently poor correlation between extension and volcanism may be partially attributed to the competing effects of thermal weakening and rheological hardening associated with intrusion and underplating of mantle-derived magmas.

Since mid-Miocene (-17 Ma) another major phase of extension, with characteristic deep-penetrating (10-15 km) block faulting and association with bimodal (basaltic- rhyolitic) volcanism, has led to formation of the Basin and Range province, where the total extension is estimated to be between 50% and 300% [Hamilton and Myers, 1966; Wernicke, 1992]. It is commonly recognized that this younger phase of basin-and-range extension [Zoback et al., 1981] is fundamentally different from the earlier low-angle detachment faults [Burchfiel et al., 1992; Coney, 1987], although the change between these two types of extension was gradual in many places [Zoback et al., 1981].

The cause of extension in the Cordillera is a subject of intensive study and debate. One view attributes the Cordilleran tectonics to plate interactions along the western margin of North America [Atwater, 1970; Severinghaus and Atwater, 1990], where subduction of the Farallon plate under North America has been replaced by the evolving San Andreas transform fault since 25-30 Ma. Sonder et al. [1987] and Wernicke et al. [1987], noting that much of the Cordilleran extension happened when western North America was under a compressive and transpressive tectonic regime, have emphasized the role of gravitational collapse. Because of the involvement of major mantle upwelling in the basin-and- range extension, there are also numerous suggestions for a causative role of mantle thermal perturbations [Parsons et at., 1994; Saltus and Thompson, 1995; Suppe et at., 1975].

a b

B A P- pgz

Prn :' ' Figure 2. Conceptual model of crustal collapse. (a) Sketch of a thickened crust with an Airy-type crustal root. (b) Lithostatic pressures along vertical profiles across the lowland crust (line A) and the mountain range (line B). The differential pressure Ap tends to cause crustal collapse; the lateral pressure gradient in the transition zone between the mountain range and the lowland tends to drive lateral extrusion of the ductile lower crust. Notice that Ap vanishes below the crustal welt.

The approach we take here is to first constrain the effects of crustal collapse of the Sevier-Laramide orogen, which involves fewer uncertainties than either plate interactions or mantle processes. By isolating the effects of crustal collapse, we may reach a better understanding of the role of mantle upwelling and plate interactions.

3. Crustal Collapse

3.1. Driving Forces and Thermal-Rheological Control

The dynamic instability of an overthickened crust at isostatic equilibrium is illustrated in Figure 2. The crustal welt is unstable because at any depth above the compensation level (taken to be at the Moho of the mountain range) the lithostatic pressure under the mountain range is greater than that under the surrounding lowland. This differential pressure tends to drive gravitational collapse of the mountain range. The vertical integral of the differential pressure represents the total extensional force [Lynch and Morgan, 1987]:

L

l[Pt(Z)-Pr(Z)],:lZ -h (])

where Pt (z) and Pr (z) are the lithostatic pressure at depth z under the mountain range and the lowland, respectively; h is the elevation of the mountain range above the reference lowland, and L is the compensation depth. The extensional force F acts on the sides perpendicular to the plane of Figure 2 and therefore has dimensions of force per unit length (N m -•) [Turcotte and Schubert, 1982]. The value of F, represented by the shaded area in Figure 2b, is numerically equivalent to the excess gravitational potential energy stored in a column of the mountain range relative to that in the lowland [Molnar and Lyon-Caen, 1988]. It is clear from Figure 2 that F is dependent on the density of the crust and mantle and the elevation of the mountain range above the reference lowland,


which may be related to crustal thickness if a uniform crustal density and the Airy isostasy can be assumed:

h = (Pm -Pc )(Ht- Hr ) Pm

(2)

where H t and H r are the thicknesses of the thickened and reference crust, respectively; Pm and Pc are the mantle and crustal density, respectively. The extensional force defined in (1) can be easily calculated from the elevation and the density contrast between the crust and the mantle [Molnar and Lyon- Caen, 1988]:

collapse, because the continental lithosphere has an• intrinsically stratified rheology (Figure 3). The strength envelopes of the lithosphere are defined by

O'yield(Z)= min(%, o a) (4)

where o't, and o' a are the stress difference ((o' 1 - o 3) / 2 ) needed for brittle and ductile extension at depth z, respectively. Laboratory experiments [Brace and Kohlstedt, 1980] suggest that ductile deformation depends on lithology, strain rate, and, especially, temperature:

h+AH / F = pcgh H r +• (3) 2

where g is gravitational acceleration and Mar is the thickness of the Airy-type crustal root: Mar = H t -H r -h. For Airy-type isostasy, Mar = hPc/(Pm-Pc)' The crust in the hinterland of the Cordilleran fold-and-thrust belts was thickened to -60 km

near the end of the Sevier-Laramide orogeny [Coney and Harms, 1984; Parrish et al., 1988]. If we take the reference crust (the crust adjacent to the hinterland) to be 40 km thick, the crustal and mantle density to be 2800 kg m -3 and 3300 kg m -3, respectively, we obtain an elevation of 3 km from (2); the corresponding extensional force is about 4.2x1012 N m -1 from (3), comparable with typical tectonic forces associated with ridge push and slab pull [Bott, 1993; Forsyth and Uyeda, 1975].

Notice that the gravitational driving forces arising from an overthickened crust alone are limited to the crust (Figure 2). This simple fact has important ramifications for crustal

where k is the strain rate; the typical value of • for continental deformation is in the range between 10 -14 s -1 and 10-16 S-1. The parameter A is a constant, H is the activation enthalpy, R is the gas constant, and T is the absolute temperature. For a given range of stresses the parameter n is a constant associated with deformation mechanism (n --3 for dislocation creep in the lithosphere). All results discussed below assume a granitic crust (A=10 -8'8 MPa -n s-1, H=123 KJ tool -1, and n=3) and an olivine-dominated mantle lithosphere (A=103'28 MPa -n s-1, H=420 KJ mo1-1, and n=3) [Kirby and Kronenberg, 1987; Korato et al., 1986]. Using other published rheological parameters for the lithosphere will not affect the general conclusions drawn here.

Brittle deformation of rocks is characterized by sliding on fractures and faults and is generally independent of strain rate, temperature, and lithology [Byerlee, 1978]:

Yield Strength (MPa) 0 100 200 300 400

0

• 20 [ Ductile ß

• 40

60

80

zSP (MPa) 0 100

Figure 3. (left) Strength envelopes of a model continental lithosphere. The dashed profile is for the reference lithosphere characterized by a 30 km crust and an equilibrium geotherm with a surface heat flux of 60 mW m -2. The shaded profiles are for 20 and 50 m.y. after an instant crustal thickening that increased the crustal thickness to 50 km; reduction of the lithospheric strength is due to postkinematic thermal relaxation and radioactive heating and the mantle lithosphere being pushed to a greater depth. Notice the thick channel of ductile lower crust under orogens. (right) The differential pressure (Ap) in Figure 2 is replotted.

t•t, =ktt• n (6)

where kt is the frictional coefficient (taken to be 0.85) and t• n is normal stress on the fault plane. Assuming fractures occur in all orientations [Brace, 1972], cr n can be replaced by the effective lithostatic stress (the lithostatic pressure minus the pore fluid pressure). Recent studies indicate that the brittle strength becomes less sensitive to pressure as depth increases [Shimada, 1993]; the complexities of t•/• at higher pressures are not critical for our discussion here and therefore are not

considered. The yield strength of the lithosphere is usually defined as the vertical integral of the yield stress across the whole lithosphere [Ranalii, 1995]:

0

S = I O'y ield (z)dz (7) l

where l is the base of the reference lithosphere. We also define the crustal strength as the yield stress integrated across the crust.

The "jelly sandwich" rheological structure shown in Figure 3 is intrinsic to the continental lithosphere because of the temperature dependence of rheology and different compositions of the crust and mantle. It is also clear from Figure 3 that the ductile lower crust is particularly weak and

LIU AND SHEN: CRUSTAL COLLAPSE AND CORDILI.ERAN EXTENSION 315

thick under orogenic belts owing to heating associated with crustal thickening (thermal relaxation, radioactive heating, shear heating, etc. [see Liu and Furlong, 1993]) and the fact that the mantle material is pushed down to a hotter regime [Glazner and Bartley, 1985]. Such rheological structures raise the important question of whether crustal collapse, driven by extensional forces arising within the crust, can lead to whole- lithosphere extension.

3.2 Crust-Mantle Decoupling and Evolution of Crustal Collapse

The effects of crustal collapse on mantle lithosphere depend on mechanical coupling between the crust and mantle, which, in turn, is influenced by thermal structures of the lithosphere. Predicting thermal structures of orogens inevitably involves poorly constrained factors such as shear heating, erosion, and heat input from the mantle [Liu and Furlong, 1993]. To derive some general constraints, we consider here two end-member cases. The first case is for a relatively cold lithosphere where the strength of the uppermost mantle prevents it from flowing together with the lower crust. This is the assumption in most models of ductile flow within the lower crust [Bird, 1991]. For crustal collapse to induce extension in the mantle lithosphere, sufficient shear stresses need to be transmitted through the ductile lower crust. The channel Poiseuille flow may be used to approximate ductile flows within the lower crust driven by the lateral pressure gradient induced by topographic changes; the shear stress exerted on the top of the mantle lithosphere is [Bird, 1991]:

z= bPcgC•x x (8)

1993]; this is a condition representative for core-complex formation in the Cordillera [Parrish et al., 1988; Wernicke et al., 1987]. The decrease of extensional force and lithospheric strength with time in Figure 4 is due to thinning of the crustal welt by crustal collapse, which in the model occurs in the form of lateral ductile extrusion within the lower crust, driven by an initial lateral topographic gradient of 0.01. The inset in Figure 4 shows the typical velocity profile. For simplicity, brittle deformation is not included; this is not critical here

because the extensional forces and lithospheric strength are mainly determined by the crustal thickness and thermal structures and are not sensitive to the details of extensional

processes. Since crustal collapse is decoupled from the mantle lithosphere, no asthenospheric upwelling is induced to compensate for the lost gravitational potential energy, so the total extensional force decreases rapidly; within 10 million years or so it becomes less than the yield strength of the crust, and crustal collapse is expected to stop. We found that, within reasonable ranges of crustal thickness (40- 60 km) and thermal structures, the predicted lifespan of crustal collapse is about 5-15 million years. Such a relatively short lifespan of crustal collapse is comparable with that of core-complex formation in the Cordillera [Parrish et al., 1988].

The other end-member case is for a relatively hot lithosphere where the uppermost mantle is sufficiently weak to flow together with the lower crust during crustal collapse; in this sense the crust and mantle are fully coupled. We have modeled ductile flow within the crust and mantle lithosphere induced by crustal thickening at orogenic belts. Figure 5a shows the model geometry and boundary conditions. The ductile flows are driven by the lateral pressure gradient

where g is the gravitational acceleration, c (equal to cos (Moho slope)) is a small geometric correction factor, and dh/dx is the topographic gradient. The parameter b is roughly half the thickness of the ductile channel for Newtonian fluids

but is only about one fourth for power law fluids [Bird, 1991]. For an upper bound stress estimation, take b = 10 km, a topographic gradient of 0.01, and a flat Moho, we find •=2.8 MPa. Similar results can be obtained from a Couette flow

approximation, assuming the lower crust flows by the shear of the sliding upper crust [Hopper and Buck, 1996]. Given the typical lithospheric strength (>1012 N m -1) [Lynch and Morgan, 1987], such shear stresses are unlikely to cause significant deformation in the mantle lithosphere. In this case, crustal collapse may be mechanically decoupled from the mantle lithosphere, as assumed in previous models [Bird, 1991; Block and Royden, 1990].

Crustal collapse that decouples from the mantle lithosphere is expected to be short-lived. This is illustrated in Figure 4, which compares the total extensional force with the yield strength of a model lithosphere determined in a one- dimensional thermomechanical model, assuming complete mechanical decoupling between the crust and mantle lithosphere. In the model the crust with an initial thickness of 35 km is thickened instantly to 55 km by overthrusting of a 20-km crustal sheet. Crustal collapse is allowed to occur when temperature at the Moho reaches 650øC by postkinematic thermal relaxation and radioactive heating [Liu and Furlong,

4.0

3.0

2.0

1.0

0.0

0

55

0.0 Velocity (mm/yr) 0.1

S 1

Crustal Collapse ---•

0 5 10 15

Time (m.y.)

Figure 4. Evolution of the total extensional force (F) and the yield strength of the lithosphere (Si) and Crust (Sc) during crustal collapse that is mechanically decoupled from the mantle lithosphere. The inset shows one typical velocity profile of ductile extrusion in the crust (brittle extension is not included in the model). Time is after the initiation of crustal collapse. Crustal collapse is expected to stop within --10 million years when the extensional force is insufficient to overcome the crustal yield strength.


Surface T = 0øC •x.

Brittle Crust u-v=0 Thickened Reference Crust Crust

Mantle

• Lithosphere •

Base of Lithosphere u = ohv / c)y = 0 T = 1300øC

Distance

b

25-

75-

0.0 0.5 1.0 1.5

Distance (100 km)

Figure 5. (a) Model geometry and boundary conditions for ductile deformation driven by the lateral pressure gradient between the crustal welt and the lowland. Only the right half of the model is solved because of the symmetry of the model. The topographic gradient, not shown in Figure 5, is calculated assuming local isostasy. The numerical mesh used for the calculations is 51x51. (b) Snapshot of the velocity field. The thick shaded line indicates the initial Moho; the thin solid line is the Moho at 12 m.y. The mantle lithosphere is fully coupled to the crust, as shown by the continuous velocity field across the crust and mantle.

between the thickened crust and the lowland. Assuming constant crustal and mantle density and local isostasy, this lateral pressure gradient is related to the change of crustal thickness:

cgP Pc ) dH (9) '•xx = P c g (1- p m dx

where P is the pressure, x is the horizontal distance, and H is the crustal thickness. To isolate the effects of crustal collapse, we impose no lateral displacement (i.e., no shortening or stretching) at the right-side boundary, which is also taken to be a far-distance boundary (calculated to a distance of 300 km) so its influence on the velocity field near the crustal welt (Figure 5b) is proven negligible. To maximize the amount of thinning of the mantle lithosphere that may be induced by

crustal flows, we chose a rigid upper boundary and a free lower boundary so that the lateral extrusion of crustal material is all balanced by upwelling of the mantle material under the crustal welt. As shown in Figure 5b, the lateral flow is channelized in the lower crust and occurs mainly in the transition zone between the thickened crust and the adjacent lowland, where the lateral pressure gradient is the highest. The vertical flow is induced by the lateral flow, as required by volume conservation. Figure 6 shows the maximum flow rate and the accumulative thinning of the mantle lithosphere under the crustal welt for one experiment that started with a thermal structure characterized by a surface heat flux of 80 mW m -2. The sharp drop of flow rate with time in Figure 6 is due to decrease of the lateral pressure gradient as the ductile flow reduces the elevation gradient; the flow essentially stops after 10-20 million years. The total amount of thinning of the mantle lithosphere is about 4 km. Within the reasonable range of model parameters, including the amount of crustal thickening (20-30 km) and the initial crustal temperature (characterized by an equilibrium surface heat flux between 60 and 90 mW m-2), the total amount of thinning of the mantle lithosphere is less than 10 km. In other words, even when the mantle is fully coupled to the crust, crustal collapse alone does not induce significant thinning of the mantle lithosphere and upwelling of the asthenosphere. The short lifespan of crustal collapse is consistent with formation of some core complexes in the Cordillera [Parrish et al., 1988]; however, crustal collapse of the Sevier-Laramide orogen seems unlikely to have been the major cause of basin-and-range extension.

4. Mantle Upwelling and Basin-and-Range Extension

There is little doubt that basin-and-range extension has involved major mantle upwelling [Stewart, 1978]. Seismic studies indicate that the lithosphere in the Great Basin is as thin as -65 km [Benz et al., 1990; Smith et al., 1989]. An abnormally hot mantle under the Basin and Range province is also indicated by the gravity [Eaton et al., 1978], high surface feat flux (-90 mW m -2) [Lachenbruch and Sass, 1978], and

10.0

.=• 7.5 7.5

5.0 5.0 Lithospheric thinnin

2.5 2.5 • .,•

Flow rate 0.0 o.o

0 lO 20 30 40 50 Time (m.y.)

Figure 6. The maximum flow rate and the accumulative thinning of the mantle lithosphere during a model crustal collapse that is mechanically coupled to the mantle lithosphere. Time is after the initiation of crustal collapse.


seismic attenuation in the upper mantle under the Basin and Range [Rornanowicz, 1979; Smith et al., 1989]. The high elevation of the Basin and Range province (~1.5 km) is largely compensated by thermally induced mass deficiency in the mantle, although some of the mass deficiency may be related to compositional heterogeneities [Humphreys and Dueker, 1994]. Although projecting the present lithospheric structure back into the geological past is difficult, significant mantle upwelling since mid-Miocene is clearly indicated by the widespread bimodal volcanism associated with basin-and- range extension [Lipman, 1980]. There was probably strong mantle upwelling just before basin-and-range extension, as indicated by the voluminous mid-Tertiary (34-17 Ma) volcanic eruption in the Great Basin. This so called ignimbrite flare-up, recorded by more than 35,000 km 3 of silicic volcanic ash flow deposited over a region of >71,000 km 2 [Best and Christiansen, 1991], requires significant upwelling of the asthenosphere to supply both the parental magmas for some of the volcanic tuff [Grunder, 1995; Johnson, 1991] and heat for the extensive crustal anatexis [Hildreth, 1981; Liu, 1996].

Such a mantle upwelling could have played a major role in driving basin-and-range extension. The mechanics of continental rifting induced by mantle upwelling have been intensively studied [Crough, 1978; Sengor and Burke, 1978] and will not be discussed. Here we examine the general gravitational instability of the lithosphere over an upwelled asthenosphere. Figure 7 shows that an elevated lithosphere isostatically supported by a buoyancy asthenosphere is dynamically unstable and tends to collapse. The situation is similar to that of an overthickened crust (see Figure 2) with two major differences: (1) With asthenospheric upwelling the gravitational extensional forces are distributed across the whole lithosphere, and (2) heat advected by the upwelling

a b

B A P=pgz

I 1

oho_ I

Upwelleld • I Asthenosphere• I I

• Z

A B

Figure 7. Conceptual model of gravitational collapse of the lithosphere due to mantle upwelling. (a) Structure of the model lithosphere. The uplifted topography is isostatically supported by thermal buoyancy forces in the upwelled asthenosphere. The depth of isostatic compensation is at the base of the reference lithosphere. (b) Pressure profiles across the reference lithosphere (line A) and the thinned lithosphere (line B).

16

12

a

40 80 120

Depth to the Upwelled Mantle (km)

b

• 50

a• 100.

150,

Yield Strength (MPa) AP(MPa) 0 50 100 0 50

Figure 8. (a) Total gravitational driving force (F) and lithospheric strength (S) as functions of mantle upwelling, calculated assuming 15 m.y. after an instant mantle upwelling to various depths, with temperature of the upwelled mantle kept constant at 1300øC. The initial thickness of the lithosphere is 150 km. (b) The strength envelope (left) and the differential pressure (right) for mantle upwelling to 60 km depth. The numbers show the values of the total driving force and yield strength ( in 1012 N m-l).

asthenosphere may significantly weaken the lithosphere. The total extensional force and the lithospheric strength depend mainly on the amount of asthenospheric upwelling (Figure 8). The results in Figure 8 are derived with a model lithosphere that is initially 150 km thick; the driving force and the lithospheric strength are calculated for 15 million years after an instant asthenospheric upwelling, comparable to the time interval between the peak mid-Tertiary volcanism and basin- and-range extension [Best and Christiansen, 1991]. Figure 8a indicates that asthenospheric upwelling to depths shallower

318 LIU AND SHEN: CRUSTAL COLLAPSE AND CORDII.LERAN EXTENSION

than 70 km depth would provide sufficient extensional forces to cause whole-lithospheric extension. Figure 8b shows the vertical distribution of the extensional forces (shown by the differential pressure between the thinned and the reference lithosphere) and lithospheric strength for asthenospheric upwelling to 60 km depth. Although uncertainties of mantle compositions under the Basin and Range province and the exact amount of basaltic magma involved in the ignimbrite flare-up make it difficult to place tight bounds on the magnitude of mantle upwelling during the peak mid-Tertiary volcanism, asthenospheric upwelling to around 60 km depth is not unreasonable. Decompressional partial melting of typical mantle materials would require mantle upwelling to less than 50 km depth [Liu and Furlong, 1992; McKenzie and Bickle, 1988], and a significant amount of basaltic magmas was likely involved in the eruption of voluminous mid- Tertiary volcanic rocks in the Great Basin [Feeley and Grunder, 1991; Grunder, 1995; Johnson, 1991]. Such a mantle upwelling would be sufficient to trigger basin-and-range extension.

The cause of mantle upwelling under the Basin and Range remains uncertain. Our results argue against crustal collapse of the Sevier-Laramide orogen being the major cause. Harry et al. [1993] have shown that gravitational collapse of a thickened crust could lead to significant mantle upwelling and lithospheric extension, but that may be attributed to the constant lateral stretching imposed in their model. Other causes include delamination of the mantle lithosphere [Bird, 1979], convective thinning [Houseman et al., 1981], mantle upwelling in a slabless window [Dickinson and Snyder, 1979] or slab gap [Severinghaus and Atwater, 1990] associated with the migration of the Mendocino triple junction, subduction- induced mantle upwelling in a back arc setting [Stewart, 1978], and a mantle plume [Parsons et al., 1994; Saltus and Thompson, 1995; Suppe et al., 1975]. Unfortunately, many of these processes are difficult to test. We may group the proposed mantle processes into two major categories according to thermal evolution within the upwelling asthenosphere: (1) passive mantle upwelling, where mantle upwells adiabatically and then cools by conduction; and (2) active mantle upwelling, where mantle upwelling is adiabatic or superadiabatic, such as in a mantle plume, and the temperature in the upwelled mantle is maintained by some kind of convective flows. The thermal histories between these

two types of mantle upwelling are significantly different so that some general constraints may be derived by comparing the predicted stability of the lithosphere with the history of basin-and-range extension in the Cordillera.

Figure 9 shows the predicted total extensional force and yield strength of a model lithosphere with an active (case A) and a passive (case B) mantle upwelling. Both cases started with an instant mantle upwelling to 60 km depth. The crust is initially 40 km thick in the extensional zone and is 35 km thick in the reference lowland. In case A, temperature was maintained to be 1300 øC at 60 km depth with an adiabatic geotherm within the upwelled asthenosphere, while in case B conductive cooling of the upwelled mantle was allowed after its initial ascension. Assuming pure shear extension of the lithosphere with a uniform strain rate of 10-15 s-l, the transient thermal structure was calculated using a one-

dimensional advection model; the change of elevation in the extensional region was calculated assuming local isostasy. The extensional force and the lithospheric strength were then calculated by integrating the differential pressure and the yield strength across the lithosphere according to (1) and (7). Figure 9 suggests that lithospheric extension may not last for more than 10 million years for case B, because thermal buoyancy forces in the upwelled mantle quickly diminish through conductive cooling during extension. The total amount of extension (vertically averaged horizontal strain) is less than 30% (equal to 10-•5 s-• x 10 m.y.) in case B. Conversely, more than 20 million years of extension can be expected for case A, with -80% extension. These values vary mainly with the amount of mantle upwelling and the initial thickness of the crust in the extensional zone and the adjacent reference lithosphere. Initial thermal structures of the lithosphere are not critical here, because they are quickly overprinted by heat advected by the upwelling asthenosphere. The initial extensional force would be higher for a greater contrast of crustal thickness between the extensional zone and

the adjacent lowland. In any case the predicted extension for active and passive mantle upwelling is significantly different. The continued extension and volcanism in the Basin and

Range province since mid-Miocene and the large total strain (50% to 300%) across the Cordillera are more consistent with an active mantle upwelling.

5. Discussion

Gravitational collapse of an overthickened crust is conceptually simple and geologically observable, a major reason for gravitational collapse to have gained much

3.0

2.t)

0o0

0 10 20

Time (m.y.) Evolution of the total extensional force (F) and Figure 9.

lithospheric strength (S) during lithospheric extension with an active (case A) and a passive (case B) mantle upwelling. In both cases, extension starts 15 m.y. after an instantaneous upwelling of the asthenosphere to 60 km depth. During lithospheric extension the upwelled asthenosphere is kept at 1300 øC at 60 km depth in case A but cools conductively in case B.


popularity in recent years. Although mantle upwelling is often involved [Dewey, 1988; England and Houseman, 1989], its relationship with crustal collapse is ambiguous. It has been shown in some models [Govers and Wortel, 1993; Harry et al., 1993] and implied in other studies [Livaccari, 1991; Molnar and Chen, 1983] that extensional collapse of the overthickened crust of the Sevier-Laramide orogen could have led to basin-and-range extension. Our results suggest that the effects of crustal collapse are more limited than previously thought. Crustal collapse of the Sevier-Laramide orogen could have accounted for much of the localized, short-lived mid-

Tertiary (>mid-Miocene) extension in the Cordillera, including formation of many metamorphic core complexes; however, it is unlikely to have been the major cause of basin- and-range extension. A different cause for basin-and-range extension is consistent with its fundamental differences with

the detachment faults associated with core-complex formations [Coney, 1987; Zoback et al., 1981].

The role of crustal collapse is limited because the extensional forces arising from an overthickened crust are limited to the crust, while the rheology of continental lithosphere is intrinsically stratified. Although any calculations of the lithospheric rheology inevitably involve uncertainties with extrapolating the laboratory-determined rheologic parameters, there are abundant geological and geophysical observations indicating the existence of a locally weak, ductile crust and mechanical decoupling between the upper crust and the mantle lithosphere (see Kirby and Kronenberg [1987] for a review). Crust-mantle decoupling is thought to be important in thin-skinned thrusting [Ranalli and Murphy, 1987] and crustal deformation at convergent orogens [Royden, 1996]. The case for crust-mantle decoupling is especially strong for crustal collapse of orogenic belts, where the ductile lower crust is particularly thick and weak. Mechanical decoupling may also occur between finer-scale rheological layers resulting from magmatism and compositional heterogeneities in the crust [Lister and Baldwin, 1993], allowing crustal collapse to occur at much lower deviatoric stresses than those required to overcome the yield strength of the whole lithosphere. This may help to explain the diffusive crustal extension in the Great Basin that spans a greater space and time than the few well-developed metamorphic core complexes [Axen et al., 1993].

Crustal collapse that decouples mechanically from the mantle lithosphere would cause thrusting near the margins of orogenic belts, as required by the volume conservation. Thrusting concurrent with extension is observed in the Andes and the Tibet Plateau [Burchfiel et al., 1992; Molnar and Lyon-Caen, 1988] and in some Neogene extensional basins in the Mediterranean [Platt and Vissers, 1989]. Evidence for thrusting at the margins of the Cordilleran hinterland coeval with core-complex extension is scarce; this may be partly attributed to the difficulties in interpreting faults from complicated tectonic overprinting in the Cordillera [Coney, 1980]. Concurrent stretching of the mantle lithosphere by processes related to plate interactions [Dokka and Ross, 1995] would also mitigate the volume problem created by crustal collapse.

Eliminating crustal collapse of the Sevier-Laramide orogen as the direct cause of mantle upwelling under the Basin and

Range province brings us a step closer to the cause of basin- and-range extension. Localized mantle upwelling under the Cordillera may have started in the Eocene or even earlier. Liu and Furlong [ 1993] find that an increased mantle heat flux was needed to account for the high crustal temperature and plutonism associated with the Eocene crustal extension and core-complex formation in the southwestern Canadian Cordillera. The mantle upwelling could have been caused by delamination of the mantle lithosphere [Bird, 1979] or convective thinning [Houseman et al., 1981]; in either case, basin-and-range extension can be regarded as having a similar origin to the mid-Tertiary extension, both resulting from the dynamic instability of a thickened lithosphere [Sonder et al., 1987]. However, the models of delamination or convective thinning are difficult to test; conversely, the southward migration of Tertiary volcanism in the Great Basin may be more easily explained as a result of the retreating Farallon plate [Best and Christiansen, 1991; Lipman, 1980]. In any case, the intensive mid-Tertiary ignimbrite flare-up indicates a strong pulse of mantle upwelling, which may have triggered basin-and-range extension by providing the excess gravitational potential energy and by thermally weakening the lithosphere [Liu, 1996].

While our discussion has been focused on crustal collapse and mantle upwelling, there is no doubt that the evolving tectonic setting in western North America has played an important role in the Cordilleran extension. The crustal collapse of the Sevier-Laramide orogen was probably triggered by the drop of compressional stresses at the end of the Laramide orogeny as a result of the reduced convergent rate between the Farallon and the North American plates [Coney, 1987]; the basin-and-range extension is closely related to the change of plate geometry at the western margin of North America, where convergence between the Farallon and North American plates has been gradually replaced by the San Andreas transform fault [Dickinson and Snyder, 1979; Zoback et al., 1981]. This change of tectonic setting may have facilitated basin-and-range extension by reducing the compressional stresses; however, it generates no major extensional forces [Sonder et al., 1986]. If the major forces driving basin-and-range extension are the thermal buoyancy forces in the upwelling mantle, our results indicate that some kinds of convective flows within the upwelling mantle were necessary to sustain the gravitational potential energy.

6. Conclusions

1. Gravitational collapse of an overthickened crust may be largely decoupled from the mantle lithosphere; crustal collapse alone cannot induce major mantle upwelling and whole-lithosphere extension. Most crustal collapse is short- lived (10-15 million years) and may have concurrent thrusting at the margins of orogens. Crustal collapse of the Sevier- Laramide orogen is consistent with the mid-Tertiary extension and formation of metamorphic core complexes in the Cordillera; however, it is unlikely to have been the major cause of basin-and-range extension.

2. The strong pulse of mantle upwelling indicated by the ignimbrite flare-up may have triggered basin-and-range extension by weakening the lithosphere and providing the excess gravitational potential energy. The cause of the mantle

320 L1U AND SHEN: CRUSTAL COLLAPSE AND CORDILLERAN EXTENSION

upwelling is uncertain, but the continuous volcanism and extension since mid-Miocene in the Basin and Range province favor an active mantle upwelling with internal convective heating.

Acknowledgments. This work was supported by the NSF grant EAR-9506460 and the ACS/PRF grant 27925-G2 administrated by the American Society of Chemistry. We thank Clem Chase for helpful discussion and P. Coney, L. Royden, J. Platt, D. Scholl, and an anonymous reviewer for their helpful review.

References

Airy, G.B., On the computation of the effect of the attraction of mountain masses, as disturbing the apparent astronomical latitude of stations in geodetic surveys, Philos. Trans. R. Soc. London, Ser. B, 145, 101-104, 1855.

Armstrong, R.L., and P. Ward, Evolving geographic patterns of Cenozoic magmatism in the North American Cordillera: The temporal and spatial association of magmatism and metamorphic core complexes, J. Geophys. Res., 96, 13,201-13,224, 1991.

Artyushkov, E.V., Stresses in the lithosphere caused by crustal thickness inhomogeneities, J. Geophys. Res., 78, 7675-7690, 1973.

Atwater, T., Implications of plate tectonics for the Cenozoic tectonic evolution of western North

America, Geol. Soc. Am. Bull., 81, 3513-3536, 1970.

Axen, G.J., W. Taylor, and J.M. Bartley, Space- time patterns and tectonic controls of Tertiary extension and magmatism in the Great Basin of the western United States, Geol. Soc. Am. Bull., 105, 56-72, 1993.

Benz, H.M., R.B. Smith, and W.D. Mooney, Crustal structure of the northwestern Basin and

Range province from the 1986 PASSCAL seismic experiment, J. Geophys. Res., 95, 21,823-21,842, 1990.

Best, M.G., and E.H. Christiansen, Limited extension during peak Tertiary volcanism, Great Basin of Nevada and Utah, J. Geophys. Res., 96, 13,509-13,528, 1991.

Bird, P., Continental delamination and the

Colorado Plateau, J. Geophys. Res., 84, 7561- 7571, 1979.

Bird, P., Lateral extension of lower crust from under high topography in the isostatic limit, J. Geophys. Res., 96, 10,275-10,286, 1991.

Block, L., and L.H. Royden, Core complex geometries and regional scale flow in the lower crust, Tectonics, 9, 557-567, 1990.

Bott, M.H.P., Modeling the plate-driven mechanism, J. Geol. Soc. London, 150, 941- 951, 1993.

Brace, W.F., Laboratory studies of stick-slip, and their application to earthquakes, Tectonophysics, 14, 189-200, 1972.

Brace, W.F., and D.L. Kohlstedt, Limits on lithospheric stress imposed by laboratory experiments, J. Geophys. Res., 85, 6248-6252, 1980.

Burchfiel, B.C., and G.A. Davis, Nature and controls of cordilleran orogenesis, western United States: Extensions of an earlier

synthesis, Am. J. Sci., 275, 363-396, 1975. Burchfiel, B.C., D.S. Cowan, and G.A. Davis,

Tectonic overview of the Cordilleran orogen in western United States, in The Geology of North America, vol. G-3, The Cordilleran Orogen: Conterminous U.S., edited by B.C. Burchfiel, M.L. Zoback, and P.W. Lipman, pp. 407-480, Geol. Soc. of Am., Boulder, Colo., 1992.

Byerlee, J.D., Friction of rocks, Pure Appl. Geophys., 116, 615-626, 1978.

Coney, P.J., Cordilleran metamorphic core complexes: An overview, Mere. Geol. Soc. Am., 153, 7-31, 1980.

Coney, P.J., The regional tectonic setting and possible causes of Cenozoic extension in the North American Cordillera, in Continental Extensional Tectonics, edited by M.P. Coward, J.F. Dewey, and P.L. Hancock, Geol. Soc. Spec. Publ., 28, 177-186, 1987.

Coney, P.J., and T.A. Harms, Cordilleran metamorphic core complexes: Cenozoic extensional relics of Mesozoic compression, Geology, 12, 550-554, 1984.

Crough, S.T., Thermal origin of mid-plate hotspot swells, Geophys. J. R. Astron. Soc., 55, 451- 469, 1978.

Dalmayer, B., and P. Molnar, Parallel thrust and normal faulting in Peru and constrains on the state of stress, Earth Planet. Sci. Lett., 55, 473- 481, 1981.

Dewey, J.F., Extensional collapse of orogens, Tectonics, 7, 1123-1139, 1988.

Dickinson, W.R., and W.S. Snyder, Geometry of subducted slabs related to the San Andreas

transform, J. Geol., 87, 609-627, 1979. Dokka, R.K., and T.M. Ross, Collapse of

southwestern North America and the evolution

of early Miocene detachment faults, metamorphic core complexes, the Sierra Nevada, orocline, and the San Andreas fault system, Geology, 23, 1075-1078, 1995.

Eaton, G.P., The Basin and Range Province: Origin and tectonic significance, Annu. Rev. Earth Planet. Sci., 10, 409-440, 1982.

Eaton, G.P., R.R. Wahl, H.J. Prostka, D.R. Mahey, and M.D. Kleinkopf, Regional gravity and tectonic patterns: Their relation to late Cenozoic epeirogeny and lateral spreading in the western Cordillera, in Cenozoic Tectonics and Regional Geophysics of the Western Cordillera, edited by R.B. Smith and G.P. Eaton, Mem. Geol. Soc. Am., 152, 51-91, 1978.

Elison, M.W., Intracontinental contraction in western North America: Continuity and episodicity, Geol. Soc. Am. Bull., 103, 1226- 1238, 1991.

England, P.C., and G.A. Houseman, Extension during continental convergence with special reference to the Tibetan plateau, J. Geophys. Res., 94, 17,561-17,597, 1989.

Feeley, T.C., and A.L. Grunder, Mantle contribution to the evolution of Middle Tertiary silicic magmatism during early stages of extension: The Egan Range volcanic complex, east-central Nevada, Contrib. Mineral. Petrol., 106, 154-169, 1991.

Forsyth, D., and S. Uyeda, On the relative importance of the driving forces of plate motion, Geophys. J. R. Astron. Soc., 43, 163- 200, 1975.

Gans, P.B., G.A. Mahood, and E. Schermer, Synextensional magmatism in the Basin and Range Province: A case study from the eastern Great Basin, Spec. Pap. Geol. Soc. Am., 233, 58 pp., 1989.

Glazner, A.F., and J.M. Bartley, Timing and tectonic setting of Tertiary low-angle normal faulting and associated magmatism in the southwestern United States, Tectonics, 3, 385- 396, 1984.

Glazner, A.F., and J.M. Bartley, Evolution of lithospheric strength after thrusting, Geology, 13, 42-45, 1985.

Govers, R., and M.J.R. Wortel, Initiation of asymmetric extension in continental lithosphere, Tectonophysics, 223, 75-96, 1993.

Grunder, A.L., Material and thermal roles of basalt in crustal magmatism: Case study from eastern Nevada, Geology, 23, 952-956, 1995.

Hamilton, W., and W.B. Myers, Cenozoic tectonics of the western United States, Rev. Geophys., 4., 509-549, 1966.

Harry, D.L., D.S. Sawyer, and W.P. Leeman, The mechanics of continental extension in western

North America: Implications for the magmatic and structural evolution of the Great Basin, Earth Planet. Sci. Lett., 117, 59-71, 1993.

Hildreth, W., Gradients in silicic magma chambers: Implications for lithospheric magmatism, J. Geophys. Res., 86, 10,153-10,192, 1981.

Hodges, K.V., and J.D. Walker, Extension in the Cretaceous Sevier orogen, North American Cordillera, Geol. Soc. Am. Bull., 104, 560-590, 1992.

Hopper, J.R., and W.R. Buck, The effects of lower crust flow on continental extension and passive margin formation, J. Geophys. Res., 101, 20,175-20,194, 1996.

Houseman, G.A., D.P. McKenzie, and P. Molnar, Convective instability of a thickened boundary layer and its relevance for the thermal evolution of continental convergent belts, J. Geophys. Res., 86, 6115-6132, 1981.

Humphreys, E.D., and K.G. Dueker, Physical state of the western U.S. upper mantle, J. Geophys. Res., 99, 9635-9650, 1994.

Johnson, C.M., Large-scale crustal formation and lithosphere modification beneath Middle to Late Cenozoic calderas and volcanic fields, western North America, J. Geophys. Res., 96, 13,485- 13,508, 1991.

Kirby, S.H., and A.K. Kronenberg, Rheology of the lithosphere: Selected topics, Rev. Geophys., 25, 1219-1244, 1987.

Korato, S.I., M.S. Paterson, and J.D. Fitzgerald, Rheology of synthetic olivine aggregates: Influence of grain size and water, J. Geophys. Res., 91, 8151-8176, 1986.

Lachenbruch, A.H., and J.H. Sass, Models of an extending lithosphere and heat flow in the Basin and Range Province, in Cenozoic Tectonics and Regional Geophysics of the Western Cordillera, edited by R.B. Smith and G.P. Eaton, Mem. Geol. Soc. Am., 152, 209- 250, 1978.

Lipman, P.W., Cenozoic volcanism in the western United States: Implication for continental tectonics, in Studies in Geophysics: Continental Tectonics, pp. 161-174, Nat. Acad. Sci., Washington, D. C., 1980.

Lister, G.S., and S.L. Baldwin, Plutonism and the origin of metamorphic core complexes, Geology, 21, 607-614, 1993.

Liu, M., Dynamic interactions between crustal shortening, extension, and magmatism in the North American Cordillera, Pure Appl. Geophys., 146, 447-467, 1996.

Liu, M., and K.P. Furlong, Cenozoic volcanism in the California Coast Ranges: Numerical solutions, J. Geophys. Res., 97, 4941-4957, 1992.

Liu, M., and K.P. Furlong, Crustal shortening and Eocene extension in the southeastern Canadian Cordillera: Some thermal and mechanical considerations, Tectonics, 12, 776-786, 1993.

Liu, M., and K.P. Furlong, Intrusion and underplating of mafic magmas: thermal- rheological effects and implications for Tertiary tectonomagmatism in the North American Cordillera, Tectonophysics, 273, 175-187, 1994.

Livaccari, R.F., Role of crustal thickening and extensional collapse in the tectonic evolution of the Sevier-Laramide orogeny, western United States, Geology, 19, 1104-1107, 1991.


Lynch, H.D., and P. Morgan, The tensile strength of the lithosphere and the localization of extension, in Continental Extensional Tectonics, edited by M.P. Coward, J.F. Dewey and P.L. Hancock, Geol. Soc. Spec. Publ., 28, 53-66, 1987.

McKenzie, D.P., Active tectonics of the Mediterranean region, Geophys. J. R. Astron. Soc., 30, 109-185, 1972.

McKenzie, D., and M.J. Bickle, The volume and composition of melt generated by extension of the lithosphere, J. Petrol., 29, 625-679, 1988.

Molnar, P., and W.-P. Chen, Focal depths and fault plane solutions of earthquakes under the Tibetan plateau, J. Geophys. Res., 88, 1180- 1196, 1983.

Molnar, P., and H. Lyon-Caen, Some simple physical aspects of the support, structure, and evolution of mountain belts, Spec. Pap. Geol. Soc. Am., 218, 179-207, 1988.

Parrish, R.R., S.D. Cart, and D.L. Parkinson, Eocene extensional tectonics and

geochronology of the southern Omineca belt, British Columbia and Washington, Tectonics, 7, 181-212, 1988.

Parsons, T., G.A. Thompson, and N. Sleep, Mantle plume influence on the Neogene uplift and extension of the U.S. western Cordillera?, Geology, 22, 83-86, 1994.

Platt, J.P., and R.L.M. Vissers, Extensional collapse of thickened continental lithosphere: A working hypothesis for the Alboran Sea and Gibraltar arc, Geology, 17, 540-543, 1989.

Ranalli, G., Rheology of the Earth, 413 pp., Chapman and Hall, New York, 1995.

Ranalii, G., and D.C. Murphy, Rheological stratification of the lithosphere, Tectonophysics, 132, 281-295, 1987.

Romanowicz, B.A., Seismic structure of the upper mantle beneath the United States by three- dimensional inversion of body wave arrival times, Geophys. J. Royal Astron. Soc., 57, 479-506, 1979.

Royden, L., Coupling and decoupling of crust and mantle in convergent orogens: Implications for strain partitioning in the crust, J. Geophys. Res., I01, 17,679-17,705, 1996.

Saltus, R.W., and G.A. Thompson, Why is it downhill from Tonopah to Las Vegas?: A case for mantle plume support of the high northern Basin and Range, Tectonics, 14, 1235-1244, 1995.

Sengor, A.M.C., and K. Burke, Relative timing of rifting and volcanism on earth and its tectonic implications, Geophys. Res. Lett., 5, 419-421, 1978.

Severinghaus, J., and T. Atwater, Cenozoic geometry and thermal state of the subducting slabs beneath western North America, in Basin

and Range Extensional Tectonics Near the Latitude of Las Vegas, Nevada, edited by B.P. Wernicke, Mere. Geol. Soc. Am., 176, 1-22, 1990.

Shimada, M., Lithosphere strength inferred from fracture strength of rocks at high confining pressures and temperatures, Tectonophysics, 217, 55-64, 1993.

Smith, R.B., W.C. Nagy, K.A. Julander, J.J. Viveiro, C.A. Baker, and D.G. Gants, Geophysical and tectonic framework of the eastern Basin and Range-Colorado Plateau- Rocky Mountain transition, in Geophysical Framework of the Continental United States, edited by L.C. Pakiser and W.D. Mooney, Mere. Geol. Soc. Am., 172, 205-233, 1989.

Sonder, L.J., P.C. England, and G.A. Houseman, Continuum calculations of continental

deformation in transcurrent environments, J. Geophys. Res., 91, 4797-4818, 1986.

Sonder, L.J., P.C. England, B. Wernicke, and R.L. Christiansen, A physical model for Cenozoic extension of western North America, in Continental Extensional Tectonics, edited by M.P. Coward, J.F. Dewey, and P.L. Hancock, Geol. Soc. Spec. Publ., 28, 187-201, 1987.

Stewart, J.H., Basin and Range structure in western North America: a review, in Cenozoic Tectonics and Regional Geophysics of the Western Cordillera, edited by R.B. Smith and G.L. Eaton, Mere. Geol. Soc. Am., 152, 1-31,1978.

Suppe, J., C. Powell, and R. Berry, Regional topography, seismicity, quaternary volcanism, and the present-day tectonics of the western United States, Am. J. Sci., 275, 397-436, 1975.

Taylor, W.J., and J.M. Bartley, Prevolcanic extensional Seaman breakaway fault and its geological implications for eastern Nevada and western Utah, Geol. Soc. Am. Bull., 104, 255- 266, 1992.

Thompson, G.A., and D.B. Burke, Regional Geophysics of the Basin and Range province, Annu. Rev. Earth Planet. Sci., 2, 213-238, 1974.

Turcotte, D.L., and G. Schubert, Geodynamics: Applications of Continuum Physics to Geological Problems, 450 pp., John Wiley, New York, 1982.

Wernicke, B.P., Cenozoic extensional tectonics of

the U.S. Cordillera, in The Geology of North America, vol. G-3, The Cordilleran Orogen: Conterminous United States, edited by B.C. Burchfiel, M.L. Zoback, and P.W. Lipman, pp. 553-581, Geol. Soc. of Am., Boulder. Colo., 1992.

Wernicke, B., R.L. Christiansen, P.C. England, and L.J. Sonder, Tectonomagmatic evolution of Cenozoic extension in the North American

Cordillera, in Continental Extensional Tectonics, edited by M.P. Coward, J.F. Dewey, and P.L. Hancock, Geol. Soc. Spec. Publ., 28, 203-221, 1987.

Wilson, J.T., Did the Atlantic close and then re- open?, Nature, 211, 676-681, 1966.

Zoback, M.D., R.E. Anderson, and G.A. Thompson, Cenozoic evolution of the state of stress and style of tectonism of the Basin and Ranges Province of the western United States, Philos. Trans. R. Soc. London, Ser. A, 300, 407-434, 1981.

M. Liu and Y. Shen, Dept. of Geological Sciences, University of Missouri, 101 Geology Building, Columbia, MO 65211. (e-mail: mian @ geosc.missouri.edu, [email protected])

(received October 16, 1997; revised January 6, 1998; accepted January 28, 1998.)

crustal collapse, mantle upwelling, and cenozoic · extension in the cordillera, including...

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