JOURNAL OF GEOPHYSICAL RESEARCH, VOL. tOO, NO. B7, PAGES 10,255-10,269, JUNE t0, 1995

Partial melting of melt metasomatized subcontinental mantle and the magma source potential of the lower lithosphere

Dennis L. Harry Department of Geology, University of Alabama, Tuscaloosa

William P. Leeman

Department of Geology and Geophysics, Rice University, Houston, Texas

Abstract. Decompression melting of peridotitic mantle is unlikely to produce the large volumes of magma erupted during the early stages of extension in continental extensional provinces such as the Great Basin of western North America. First, a substantial amount of extension is required before asthenospheric mantle ascends to depths shallow enough to begin melting, even if it is unusually hot. The lithospheric mantle is too cold to undergo melting, even after large amounts of extension. Second, decompression melting of peridotite produces progressively greater amounts of magma during extension; in the Great Basin silicic magmatism was voluminous during the early stages of extension but diminished in volume as extension progressed. Comparison of empirical melting relations with modeled pressure and temperature conditions in the lithosphere during extension indicates that large volumes of magma may be produced by partial melting of melt metasomatized subcontinental mantle (SCM) during the earliest stages of extension. Melt-metasomatized SCM is expected to contain components less mafic than peridotite, emplaced during older magmatic episodes. Mafic components within the SCM will begin to melt during the earliest stages of extension if the base of the lithosphere is -1300øC or warmer and at depths greater than 75 km. Interaction of these melts with crustal rocks, fractionation processes, and crustal anatexis driven by the heat contained in the ascending mantle melts can produce the silicic to intermediate compositions observed in Great Basin Middle Tertiary magmas. A suite of models comparing different parameters show that the potential for the lower portions of the SCM to produce melt during the early stages of extension depends most strongly on the pre-extension thickness of the lithosphere. If the lithosphere is thicker than about 150 km, the base of the SCM will lie well within the subsolidus field for mafic rocks and a substantial amount of extension is required before it begins to melt. At pressures less than about 2.0 GPa, ascent paths in the SCM are nearly parallel to the basalt solidus, and so little melting occurs if the lithosphere is thinner than about 60 km. Optimal conditions for producing large volumes of melt during the early stages of extension require an initial lithosphere -100-150 km thick. Under these conditions the lower 25 km of the SCM may undergo partial melting during extension. Predicted magma flux rates match the space-time pattern of Middle Tertiary silicic magmatism in the Great Basin; the model predicts a high rate of melt production at the onset of extension which peaks within the first 10-20 m.y. and then decreases rapidly with continuing extension. Melt production within the SCM is consistent with trace element and isotope characteristics of mid- Miocene and younger basalts in the Great Basin, which suggest a transition at 5 Ma from old, geochemically evolved magma sources within the lithosphere to later asthenospheric sources. The later asthenospheric melts are compatible with decompression melting of peridotitic mantle after >50% extension.

Introduction

Many aspects of extension-related magmafism can be attrib- uted to decompression melting of peridotitic sublithospheric mantle during its ascent beneath the extending lithosphere. The decompression melting model is consistent with the timing and volume of basaltic eruptions in highly extended regions and with

Copyright 1995 by the American Geophysical Union.

Paper number 94JB03065. 0148-0227/95/94JB-03065 $05.00

the association of many rifted continental margins with large ig- neous provinces [McKenzie and Bickle, 1988; White and McKenzie, 1989]. The model is particularly successful in pre-

dicting the relationship between extension and magmatism dur- ing the late stages of extension and in regions thought to be un- derlain by anomalously hot mantle [Watson and McKenzie, 1991; Harry and Sawyer, 1992a, b; Arndt and Christensen, 1992].

Physical and geochemical arguments suggest that magmas emplaced during the early stages of extension are unlikely to re- sult from decompression melting of a sublithospheric mantle source [Leeman and Harry, 1993]. First, even unusually hot

10,255

10,256 HARRY AND LEEMAN: PARTIAL MELTING OF SUBCONTINENTAL MANTLE

peridotitic mantle must ascend to at least -100 km before it can begin melting (Figure 1). Unless the lithosphere is unusually thin to begin with, this requires from 50 to 100% extension. In many extended regions such as the Great Basin of western North America, large volume silicic to intermediate magmatism gen- erally is associated with the earliest extensional structures [Gans et al., 1989; Armstrong and Ward; 1991], although both mag- matism and extension are not uniformly distributed throughout the region [Axen et al., 1993; Best and Christiansen, 1991]. These magmas are probably derived predominantly by crustal anatexis, but limited petrologic work [Feeley and Grunder, 1991 ] suggests that coeval mafic magmas provided the necessary heat in which case melting conditions must have been reached at mantle depths at about the time of initial extension. Also, the amount of mafic magma required is significant and must be at least comparable to the volume of silicic magma produced. Sec- ond, silicic magmatism associated with the early stages of exten- sion typically decreases in volume with time. For example, sil- icic magmatism in the Great Basin was much more voluminous

1800

-•- 1700

::x 1500

• 1400

• 1300

- a)

Solidus

.. _1480øC .....

- ...- .- .- _ .- .- .1280øC

1200

1100 i I i i

50 1 O0 150 200 2:

Depth (km)

30

20

10

b) i i i i

1480 ø

1380oC --•'•'•'•

0 1 2 3 4 5

p Figure 1. Melting relations within the subcontinental mantle during extension [McKenzie and Bickle, 1988]. (a) Adiabatic ascent paths for mantle rocks with potential temperatures between 1280øC and 1480øC cross the dry peridotite solidus at depths between 50 and 110 kin, depending upon the potential temperature of the mantle. (b) The thickness of the melt column versus extension factor [5. [5 = (L+AL)/L, where zXL is the change in the width of the region undergoing extension and L is the initial width of the extending region.

during the Oligocene (early stages of extension) than at later times, in spite of continuing extension [Best and Christenson, 1991; Armstrong and Ward; 1991; Ward, 1991 ]. In contrast, de- compression melting of sublithospheric mantle should produce increasing volumes of magma with continued extension (Figure 1). Finally, trace element geochemistry and isotopic composi- tions of basalts erupted in the Great Basin prior to the late Plio- cene (-5 Ma) indicate derivation from old, isotopically evolved sources that are distinct from the sublithospheric magma sources that produce mid-ocean ridge basalts (MORB) and ocean island basalts (OIB). Unfortunately, the rarity of mafic lavas older than 15-20 Ma precludes direct assessment of the depth of melting or the nature of the mantle source involved at the onset of extension. However, the mid-Tertiary basalts erupted (predominantly Miocene) have low relative abundances of high field strength elements (HFSE), low 143Nd/144Nd and high 2ø7pb/2ø4pb and 87Sr/86Sr ratios [Leemah and Fitton, 1989; Fitton et al., 1991; Kempton et al., 1991; Daley and De Paolo, 1992]. In contrast, younger (<5 Ma) basalts in the Great Basin com- monly have compositions similar to OIB, with high Nb and TiO 2 contents, high 143Nd/144Nd, and low 87Sr/86Sr ratios [Fitton et al., 1991; Kempton et al., 1991 ]. The appearance of OIB-like basalts in the Great Basin is diachronous [e.g., Ormerod et al., 1985] and seems to have progressed from the south to as far north as north central Nevada [Lum et al., 1989]. The distribution of these latter lavas may be related to the onset of decompressional melting of upwelling sublithospheric mantle [Harry et al., 1993], following a critical amount of extension. The degree of extension varies regionally, but generally diminishes northward [Leeman and Harry, 1993]. Various workers have interpreted the geochemical data to indicate a transition at 5 Ma from an enriched source within the lithosphere to a source within the upwelling asthenosphere [Ormerod et al., 1985; Fitton et al., 1988; Lum et al., 1989; Daley and De Paolo, 1992; Fuerbach et al., 1993; Hawkesworth and Gallagher, 1993]. This transition from a shallow to deeper source with increasing extension is op- posite the trend expected from melting within the astheno- spheric mantle, in which case increasingly greater proportions of melt are predicted from shallower regions as the asthenosphere rises.

Leeman and Harry [1993] examined the potential for melting within the subcontinental mantle (SCM) during extension and showed that partial melting of mafic compositions entrained in otherwise infertile mantle can account for the onset and duration

of silicic magmatism and Miocene basaltic volcanism in the Great Basin. In this context, SCM is considered to be the por- tion of the upper mantle which is not involved in deep mantle convection. Mafic components may be introduced into the SCM during previous magmatic episodes, taking the form of plutonic intrusions, dikes, sills, or quenched intergranular melts [Singer et al., 1989; Lister, 1991; Spera, 1987]. In the Great Basin, widespread mafic melt metasomatism of the SCM may have re- sulted from magmatism associated with Mesozoic to Mid-Terti- ary subduction and/or Proterozoic continental crust formation. If geochemically evolved domains (i.e., enriched in radiogenic iso- topes) in the SCM are produced by antecedant melt metasoma- tism, then preservation of such heterogeneities requires that temperatures decrease to subsolidus conditions and remain so until the next magmatic episode. Any thermal perturbations or lithospheric thinning would result in partial melting of the geo- chemically anomalous domains and destruction of the geochemi- cal anomalies, leaving a residue of refractory mantle (peridotite?) with compositional features similar to the original

HARRY AND LEEMAN: PARTIAL MELTING OF SUBCONTINENTAL MANTLE 10,257

mantle (i.e., MORB-like). Alternatively, the lower SCM may become detached from the overlying plate and reincorporated into the asthenosphere [e.g., Bird, 1978; Nelson, 1992], leaving only a relatively thin layer in which geochemically evolved do- mains are preserved. Although subsequent cooling may result in an increase in the thickness of the mechanically strong litho- sphere, mantle material accreted into the SCM in this way would consist of refractory mantle which is unlikely to melt during subsequent extensional episodes.

The major results of Leeman and Harry [1993] are as fol- lows: (1) melting of mafic components in the lowermost portion of the SCM may begin immediately after the onset of extension if the lithosphere is initially -125 km thick, (2) only the lower - 25 km of the SCM is likely to undergo partial melting, even af- ter large amounts of extension, and (3) the amount of melt pro- duced in the lower SCM decreases rapidly after 25-30% exten- sion. The amount of extension required to produce melt and the duration of melting are strongly dependent upon the initial thickness of the lithosphere and the degree of cooling in the SCM during extension. In this paper, we examine the effects which the initial thickness of the lithosphere and the thermal boundary condition at the base of the lithosphere have on melt generation. Specifically, we examine the range of conditions under which melting is likely to occur and evaluate the factors controlling the timing, volume, and duration of melting. The purpose of the study is to test the hypothesis that partial melting of mafic components in the SCM may occur during the early stages of extension and to determine the range of conditions un- der which it is likely to contribute substantially to early synex- tensional magmatism. A finite difference model is used to cal- culate the pressure and temperature conditions within the litho- sphere during extension. The modeled P-T paths of material in the lithospheric mantle are compared to empirical melting rela- tions to determine the onset and amount of partial melting. The results confirm that the thickness of the lithosphere prior to ex- tension is the primary control on the potential for partial melting of mafic components in the lithosphere to produce early synex- tensional magmas.

Distinctions Between Subcontinental Mantle, Lithospheric Mantle, and Boundary Layer Models

The distinction between melting a chemically enriched or depleted source is relatively well defined on the basis of lava compositions. The distinctions between subcontinental mantle, lithospheric mantle, and asthenospheric mantle are more con- tentious. For the purposes of this paper, we regard the subcon- tinental mantle (SCM) to be the uppermost portion of the mantle that is not currently involved in asthenospheric convection. Thus the SCM (as used in this paper) includes the mechanical boundary layer and thermal boundary layer of McKenzie and Bickle [1988] (Figure 2). We regard the lithospheric mantle to be the upper portion of the mantle which (1) moves with the tec- tonic plate and is involved in extensional deformation; (2) is characterized by conduction as the dominant heat transport mechanism; and (3) remains mechanically isolated from deeper mantle convection systems for periods of the order 108 to 109 years. The lithospheric mantle may not include all of the SCM and represents a conductive thermal boundary layer which is strong enough to remain attached to the Earth's upper mechani- cal boundary layer for long periods of time (Figure 2). This

1600

solidus

Convecting Mantle

50 100 150

!

0 0 200 250

Depth (km)

Figure 2. Boundary layer model of the lithosphere [McKenzie and Bickle, 1988]. The mechanical boundary layer is the upper -100 km of Earth which is characterized by conductive heat transport and is able to support surface loads for geologically long periods of time. At depths greater than -180 km, the mantle is characterized by convective heat transport and is mechanically weak and compositionally well mixed (the asthenosphere). Between-100 and 180 km, the mantle heat transport mechanism is transitional between conductive and convective and may change with tectonic conditions. The lithosphere (stippled pattern) is taken to be the portion of the Earth which is sufficiently strong to move as a coherent tectonic plate and may include a portion of the transient thermal boundary layer. This portion of Earth may remain isolated from deep mantle convection for periods of 108-109 years, forming a chemical boundary layer with isotopic and trace element characteristics distinct from the underlying asthenosphere.

portion of the mantle is free to evolve trace element and isotope concentrations which are distinct from the convecting mantle, resulting in more radiogenic Sr and Pb and less radiogenic Nd isotopic ratios relative to the deeper convecting mantle regions. Mafic lavas with such isotopic compositions commonly are in- ferred to be derived from a lithospheric source [White, 1988; Gallagher and Hawkesworth, 1992; Hawkesworth and Gallagher, 1993]. These geochemical characteristics appear to require isolation from asthenospheric convection for periods of- 1 b.y. or greater [DePaolo, 1988; Kempton et al., 1991; Gallagher and Hawkesworth, 1992; Smith, 1993]. We will adopt the simplifying assumption that the compositionally dis- tinct mantle extends to the base of the conductive thermal

boundary layer (Figure 2), typically given by the depth of the 1333øC isotherm [Parsons and McKenzie, 1978]. Material above this depth is assumed to lie within the lithosphere and remain mechanically and compositionally distinct from the un- derlying convective asthenosphere between tectonomagmatic events [e.g., Hawkesworth and Gallagher, 1993].

The conductive thermal boundary layer may be underlain by a transient thermal boundary layer, which is characterized by con- ductive heat transport during periods of quiescence and vigorous convection during periods of tectonism [Parsons and McKenzie, 1978; Jordan, 1988]. Portions of the transient thermal boundary layer may also be isolated from deeper mantle convection and so represent the lower part of the SCM. The transient thermal boundary layer may become compositionally distinct from the underlying convecting mantle if long periods of quiescence al-

10,258 HARRY AND LEEMAN: PARTIAL MELTING OF SUBCONTINENTAL MANTLE

low it to remain isolated from asthenosphere convection for long periods of time. However, it is not regarded as part of the li- thospheric mantle unless it is strong enough to move coherently with the tectonic plate and become extensionally deformed.

Anderson [1994] argues that olivine-rich mantle above 650- 750øC is too weak to remain attached to the lithosphere for sub- stantial periods of time. If the 700øC isotherm is taken as the lower limit of the mechanical boundary layer (about 70 km in Figure 2), then it is unlikely that basaltic magmas (with liquidus temperatures >1100øC) will ever be produced within the li- thospheric mantle. Anderson proposes instead that isotopically evolved basaltic magmas are derived from a more or less worldwide metasomatized, buoyant "perisphere" layer situated below the lithosphere and above strongly convecting astheno- sphere. Beneath continents, this would represent the lower SCM referred to above. This model implies that there should be little or no correlation between magma isotopic chemistry and characteristics of the tectonic plates. However, in the western United States there seems to be a significant correlation between basalt isotope chemistry and age of the underlying crust/lithosphere [Leeman, 1982; Menzies, 1989; Leeman et al., 1992]. This implies that magmatic chemistry is at least partly inherited from sources entrained within the lithosphere. Ques- tions remain concerning the depth to which compositional anomalies are preserved (i.e., what is the long-term mechanical thickness of the lithosphere and under what conditions will such lithosphere produce magma?).

A Model for Melting of Metasomatized Subcontinental Mantle

The model used in this study assumes that the lithosphere de- forms by pure shear and that lateral heat conduction can be ne- glected [McKenzie, 1978]. The simple shear model is adopted in order to examine the uplift and cooling history of the mantle at any one location within an extending region. In reality, ex- tension is a complicated process and may be nonuniformly dis- tributed between the crust and mantle, the loci of extension in the crust and mantle may differ, and their relative positions may vary dynamically with time [Lister et al., 1986; Braun and Beaumont, 1987; Dunbar and Sawyer, 1989; Harry and Sawyer, 1992b]. For example, mantle strain may be strongly focused near structural features (e.g., sutures), resulting in a relatively large amount of localized mantle upwelling within a province undergoing diffuse crustal extension. In such instances, the amount of magma produced by decompression melting within the lithosphere will be locally more voluminous but less widely distributed than predicted by the simple shear model. Nonethe- less, at any given location the lithospheric mantle will follow a decompression and cooling history which places it at lower tem- peratures and pressures than the underlying asthenospheric mantle at any given time. As will be shown below, for a wide range of initial lithosphere thickness and thermal structure this requires the lithospheric mantle to cross into the basalt super- solidus field before asthenospheric mantle crosses into the dry peridotire supersolidus field. The mafic components in the li- thospheric mantle will then begin to melt prior to peridotitic mantle, regardless of the degree to which strain is focused or partitioned. Thus, although focusing of strain and vertical strain partitioning may complicate the relative timing and spatial dis- tribution of melt production in the lithosphere and mantle, the simple shear model is a valid indicator of the melting sequence.

Similarly, although arguments have been presented for en- hanced melting of peridotitic mantle as a result of small-scale convection in the sublithospheric mantle during extension [e.g., Mutter et al., 1988; Zehnder et al., 1990], such melting will also usually postdate melting of mafic material within the li- thospheric mantle. Small-scale convection serves primarily to circulate large volumes of asthenospheric mantle to depths shallower than the peridotite solidus. This results in a larger volume of melt produced in the asthenospheric mantle after a given amount of extension than predicted in the absence of small-scale convection. However, no melt is produced in the as- thenospheric mantle (with or without small-scale convection) until a sufficient amount of extension has occurred to allow as-

thenospheric material to ascend to supersolidus pressures. The model presented here demonstrates that melting of mafic mate- rial in the lithospheric mantle will often occur well before the lithosphere has thinned sufficiently to allow asthenospheric mantle to ascend to depths shallow enough to begin melting. The model shown below thus addresses the problem of produc- ing significant volumes of melt during the earliest stages of ex- tension, before asthenospheric mantle has ascended sufficiently to begin melting.

If cross-sectional area is conserved during extension, the depth to the base of the lithosphere through time is given by

Za(t) - Zø (1) (l + Uxt/Lo)

where Z 0 and L 0 are the initial depth to the base of the litho- sphere and width of the extending region, respectively, and U x is the net rate of extension (measured across the entire width of the extending region; Figure 3a). The rate of lithospheric thin- ning decreases with time as the region undergoing extension widens (Figure 3b).

Thinning is uniformly distributed in the simple shear model, so equation (1) can be generalized to give the ascent path of ma- terial throughout the lithosphere:

Z(z',t) = Z a (t)(1 - z') z' = (Z o - z)/Z o . (2)

The dimensionless variable z' identifies an individual packet of rock, initially at depth z, and equation (2) tracks its ascent (measured relative to the surface) through time.

The temperature within the lithosphere is given by the one- dimensional heat equation. In the nondimensional reference frame [ z', t }, this is

pC aT- A(z') + K a2T +pcuz(z',t)Ga (3) dt dz '2

where T(z',t) is the temperature of each rock packet, A(z5 is the heat production, K is the thermal conductivity, U•(z', t) is the rate at which a packet of rock is ascending, and G a is the adi- abatic temperature gradient in the mantle (the appendix). The thermal parameters are given in Table 1. We assume an expo- nential decay in heat production within the crust and no heat production within the mantle [Lachenbruch and Sass, 1977; Morgan, 1984; Sclater et al., 1981]:

A(z) = Ao e-z/ø 0 < z < z m (4)

A(z) = 0 Z m < Z < Z a.

Equation (3) is solved using an explicit second-order finite difference approximation in space and the fourth-order Runge-

HARRY AND LEEMAN: PARTIAL MELTING OF SUBCONTINENTAL MANTLE 10,259

U o

! Lo ! a) r(0, t)=0 T=0

T = T a or Q = Qo

T(Za,t ) = T• (7)

car[ =

where Zm + and z m' indicate evaluation of the third condition just above and below the base of the crust. Equations (6) and (7) have the solution [Turcotte and Schubert, 1982]

T(z) = AøD2 Kc (1 - e -z/ø) + qm Z 0 < z < z m /Cc - (8)

-[ Ho Uz

' L o + U x ß t [

1.4

• •.0

• 0.8

• 0.6

•,• 0.4

b)

I I I I

0.20 20 40 60 80 100

Time (m. y.)

Figure 3. Simple shear model of extension in the lithosphere. (a) A region of lithosphere initially of width L o and thickness H o extends at a rate U•. After time t, the extended lithosphere has a width of L 0 + Uff and a thickness Ho-Uzt. (b) The rate of thinning U z decreases with time.

Kutta method to step the solution forward in time [Dahlquist and Bjorck, 1974]. Pressure is determined by integrating den- sity over depth:

z •

e(z') = gJ p(z)az (5) 0

where g is the acceleration of gravity. Hence the model tracks the P-T conditions of individual packets of rock through time during extension.

The temperature within the model prior to extension is de- termined from the one-dimensional steady state heat equation:

d2T = -•1 A(z). (6) dz • K

Specifying the temperature at the surface and base of the model and requiring the heat flux from the top of the mantle to equal the heat flux into the base of the crust leads to the boundary conditions

T(z) = T•qm (z- Z a) z m < z < z a. (9) The initial thermal structure of the lithosphere is thus deter- mined by the thermal conductivity of the crust and mantle, K c and K m, the parameters describing the distribution of heat pro- ducing elements in the crust, D and A 0, the depth to the base of the crust and lithospheric mantle, z m and %, and the temperature and heat flux at the base of the lithosphere, T• and qm'

The temperature in the model lithosphere, and hence its melting behavior, depends strongly upon the nature of the boundary condition at the base of the model. End-members are the constant temperature and constant heat flux boundary condi- tions. In the former, the base of the lithosphere is prevented from cooling during extension, resulting in a relatively hot li- thospheric mantle and promoting melting in the lower litho- sphere. Figure 4a shows the pressure and temperature condi- tions predicted within the lithospheric mantle for a model in which the lithosphere is initially 125 km thick and the base of the lithosphere is held at 1333øC during extension. The initial width of the region undergoing extension is 500 km, and the ex- tension rate is 5 mrn/yr. Other model parameters are listed in Table 2. The subvertical dashed lines indicate the P-T path of a given packet of rock through time. For example, the base of the lithosphere (which is held at a constant temperature of 1333øC during extension) follows an isothermal path of decompression from about 3.8 GPa to 1.7 GPa. The crust/mantle boundary is initially at 1.2 GPa (40 km) and 620øC and follows a curved path of ascent as it decompresses and cools during extension. The subhorizontal solid lines in Figure 4a indicate the range of P-T conditions between the base of the crust and the base of the

lithosphere after given amounts of extension. For example, prior to extension the P-T conditions within the lithospheric

Table 1. Thermal Parameters

Parameter Value

A0 surface heat production rate D heat production decay constant Ga adiabatic temperature gradient Kc thermal conductivity (crust) Km thermal conductivity (mantle) Pc density at OøC (crust) P,n density at OøC (mantle) Cc specific heat (crust) Cm specific heat (mantle) •c coefficient of thermal expansion (crust) •rn coefficient of thermal expansion (mantle)

2.5 x 10 -6 W m -3

10km

0.6 øC/km

2.5 W m -1 øC-1

3.4 W m -1 øC-1

2.9x 103 kg m -3 3.3x103 kg m -3 875 J øC-1 kg 4 1250 J øC-! kg 4 2x 10 '5 øC-1 2.4x 10-6 øC-1

10,260 HARRY AND LEEMAN: PARTIAL MELTING OF SUBCONTINENTAL MANTLE

4.0

3.0

1.0

a)

I Constant Temperature Boundary Condition U x = 5 mrn/yr Zo= 125kin

crust/mantle

boundary

50

100

base of lithosphere

125 km

120/I 0

110/ •

I 20

100 m.y.

4OO

4.0

3.0

¸

1.0

b) Constant Heat Flux Boundary Condition U x = 5 mm/yr Zo= 125 km

base of lithosphere

125 km

120 10

11o

•oo/•'

90 / •

• 20

crust/mantle

boundary

70 /

60

.0 I I 600 800 1000 1200 1400 1600 400 600 1400

o

Temperc]ture C

800 1000 1200 o

Temperc:ture C

Figure 4. Pressure versus temperature diagrams for the lithospheric mantle during extension, assuming an extension rate of 5 mm/yr. The dark lines indicate dry tholeiite and peridotite solidi. The light solid lines show the range of P-T conditions in the mantle at the indicated times after the onset of extension. The light dashed lines indicate the P-T paths of individual rock packets through time. Model parameters are those for model 1 in Table 2. (a) Constant temperature lower thermal boundary condition (T a = 1333øC). (b) Constant heat flux lower thermal boundary condition (qm = 31 mW/m2). (c) Adiabatic lower thermal boundary condition.

i

1600

mantle range from about 1.2 GPa and 620øC to 3.8 GPa and 1333øC (uppermost solid line, Figure 4a). After 100% exten- sion, the P-T conditions range from 0.6 GPa and 560øC to 1.9 GPa and 1333øC (lowermost solid line, Figure 4a).

No portion of the model lithosphere ever crosses the dry peri- dotite solidus, so dry peridotitic lithospheric mantle is unlikely to melt during extension. If the mantle were partially saturated with volatiles, melting of peridotitic mantle during the early stages of extension is possible [e.g., Gallagher and Hawkesworth, 1992]. However, volatile components will be preferentially incorporated into the earliest melts [Leeman and Harry, 1993]. Continued melt production over the -10-15 m.y. duration of early voluminous volcanism in the Great Basin seems unlikely unless a mechanism exists to continuously re- plenish the volatiles. Although subduction processes have been

suggested as a mechanism for introducing volatiles into the li- thospheric mantle [e.g., Gallagher and Hawkesworth, 1992; Smith, 1993], the volume of magma inferred to be produced in this way in modem magmatic arcs is much less than that em- placed in the Great Basin following the onset of extension. Thus, although magma production from volatile enriched mantle may contribute to early synextensional magmatism in the Great Basin, it appears that wet peridotitic mantle is unlikely to sus- tain voluminous melting for long periods of time.

An alternative scenario involves melting of mafic components entrained within the lithospheric mantle. If the base of the lithosphere does not cool rapidly during extension, the lower- most lithosphere will cross the basalt solidus during the earliest stages of extension (Figure 4a). Partial melting of mafic com- ponents in the lithospheric mantle may then result in early syn-

Table 2. Model parameters

Parameter

qm

qs

Model 1 Model 2 Model 3

depth to base of crust, km depth to base of lithosphere, km temperature at base of model, øC temperature at base of crust, øC mantle heat flux, mW m -2 surface Heat Flux, mW m -2

40 40 40

125 75 175

1333 1333 1333

578 851 455

31 46 22

52 68 44

HARRY AND LEEMAN: PARTIAL MELTING OF SUBCONTINENTAL MANTLE 10,261

4.0 -

3.0

c)

1.0

Adiabatic Boundary Condition U x = 5 mm/yr Zo= 125 km

crust/mantle

boundary 60

50

lOO

9o/' /( /l

/ I//

70/ •

base of lithosphere

125 km

/ •

I

/,

m.y.

.0 • • • • •

400 600 800 1000 1200 1400 1600 o

Temper(:ture C

Figure 4. (continued)

extensional magma production. Such a scenario may be appro- priate in areas characterized by relatively high extension rates, where advective transport of heat in the rising mantle outpaces conductive cooling.

Figure 4b shows a similar model, in which the heat flux at the base of the lithosphere is maintained at the initial value of 30 mW/m 2 during extension. The base of the model lithosphere cools by about 300øC after 100 m.y. (100% extension). With this amount of cooling, no portion of the lithosphere or underly- ing mantle is ever likely to melt. Such a scenario may be appro- priate for regions undergoing slow rates of extension such as the Viking Graben [Pealerson and Ro, 1992; Latin et al., 1990], which are typically associated with little early synextensional magmafism.

The fact that magmatism ultimately develops in areas sub- jected to more than-100% extension suggests that the constant heat flux model (Figure 4b) is not an adequate description of the lithosphere's thermal state in many large extensional provinces such as the Great Basin. An alternative boundary condition is shown in the model in Figure 4c. In this model, the astheno- sphere is assumed to rise adiabatically [McKenzie and Bickle, 1988]. For continuity of temperature at the lithosphere-astheno- sphere boundary, the base of the model is specified to follow the adiabatic cooling path for material beginning at 1333øC and 125 km depth, which results in moderate cooling at the base of the lithosphere (Figure 4a). Both the constant temperature bound- ary condition and the adiabatic boundary condition produce models which predict partial melting of mafic material in the lower lithosphere during the earliest stages of extension. How- ever, the moderate cooling of the base of the lithosphere in the adiabatic model (about 30øC during the first 100% of extension) slightly diminishes the thickness of the zone of melt production.

The large volume of basalt emplaced in highly extended prov- inces indicates that mantle as warm as -1280øC must rise to

shallow depths during extension [e.g., McKenzie and Bickle, 1988]. Therefore the models shown in Figures 4a and 4c are considered more realistic in highly extended regions than the constant heat flux model (Figure 4b). We will take the adiabatic lower thermal boundary condition as a first-order approximation of the lithosphere's behavior at extension rates between -2.5 and 10 mm/yr. Models using the adiabatic lower thermal boundary condition produce similar amounts of melt after a given amount of extension at all rates of extension between these extremes,

indicating the model is relatively insensitive to modest vari- ations in extension rate (Figure 5).

Melting Relations in the Subcontinental Mantle

As discussed above, melting of dry peridotitic mantle re- quires that heat is added to the lithosphere, raising its tempera- ture. Alternatively, the solidus temperature of the mantle may be lowered (e.g., by the presence of volatile components). If volatile-bearing phases are present [e.g., Gallagher and Hawkesworth, 1992], much of the lower lithosphere may lie within the supersolidus field prior to extension (Figure 6). Be- cause the melt is buoyant, this portion of the lithosphere will presumably already have had any melt extracted and so will be depleted of easily fusible components. Synextensional melting of wet mantle only occurs if packets of rock move further into the supersolidus field or cross from the subsolidus field into the supersolidus field during extension. At pressures less than about 2.0-2.5 GPa (shallower than -60-75 km), the melting temperature of fluid saturated pyrolite either increases upward or remains fairly constant. Hence little melting of fluid-enriched mantle is predicted below these pressures. For example, the dashed line representing the packet of rock initially at 80 krn in Figure 6 lies near the water-saturated pyrolite solidus prior to extension but never moves far into the supersolidus field during extension. Melting of this packet of rock ceases at pressures less than -1.7 GPa, as the melting temperature of water-satu- rated pyrolite begins to increase. Portions of the lithospheric mantle initially at pressures greater than about 2.5 GPa may move further into the supersolidus field during ascent and pro- duce early syn-extensional melting. As stated previously, though, unless a mechanism is available to replenish the vola- tiles they will be rapidly depleted by the earliest melts, resulting in a shift toward the dry peridotire melting relation and short- lived magmatism. Furthermore, volatile-saturated conditions such as those shown in Figure 6 are unlikely to be representative of the mantle. More realistic fluid-undersaturated conditions

will result in solidi lying closer to the dry peridotire solidus shown in Figure 6, thus restricting both the degree of volatile- enhanced partial melting and the depth range over which melt- ing can occur.

Because neither dry or fluid-enriched melting of peridotire mantle appears adequate to explain early synextensional mag- mas in extensional provinces such as the Great Basin, we con- sider melting of mafic components in the lithospheric mantle. As Figure 6 shows, if the lithosphere is initially -125 km thick [Leeman and Harry, 1993], mafic components begin to melt during the earliest stages of extension. In the next section we discuss the effect which the initial thickness of the lithosphere has on the timing and duration of melt production in mafic com- ponents of the lower lithosphere. As discussed previously, ma-

10,262 HARRY AND LEEMAN: PARTIAL MELTING OF SUBCONTINENTAL MANTLE

4.0

3.0

1.0

4.0

3.0

a)

I Adiabatic Boundary Condition U x = 2.5 mm/yr Zo= 125 km

90

80

70: crust/mantle

/

boundary 60/

100

400 600

base of lithosphere

125 km

120 10

11o

800 1000 1200 o

Temperature C

% extension

(200 m.y.)

i

1400

b)

I Adiabatic Boundary Condition U x = 10 mm/yr Zo= 125 km

base of lithosphere

125 km

120/I 0 110/ i •

90

80

crust/mantle

boundary

50

70 /

% extension

60 (200 m.y.)

I

1600

1.0

400 600 800 1000 1200 1400 1600 o

Temperature C

Figure 5. Effect of extension rate on model results. (a) Model 1, with extension rate U• = 2.5 mm/yr. (b) Model 1, with extension rate U• = 10 mm/yr. The model results are relatively insensitive to variations in extension rate within this range.

tic phases may be present in highly metasomatised mantle as remnants of past magmatic episodes in which sublithospheric melts intruded into the shallow mantle [Pegram, 1990; Lister, 1991]. A portion of these mafic magmas would be quenched during their ascent, becoming stranded in the lithospheric man- tle [Singer et al., 1989; Spera, 1987].

Results

Figure 7 shows the results for models with initial lithospheric thicknesses of 75 and 175 km and an adiabatic lower thermal

boundary condition (compare to Figure 6). In the case of an in- itially thin lithosphere (Figure 7a), all of the lithospheric mantle initially deeper than about 62 km lies in the basalt supersolidus field prior to extension, and remains in the supersolidus field throughout the extensional episode. Presumably, most of the melt present prior to extension would have ascended, leaving a more refractory bulk composition. During extension, little addi- tional melt is produced because no portion of the mantle moves much further into the supersolidus field than it was to begin with. After 100 m.y. (100% extension), only an additional 2 km of material has crossed from the subsolidus into the supersolidus field (the material initially between 60 and 62 km). Most melt- ing occurs during the first 30-40% of extension, at which point material initially at 60 km depth is just approaching the solidus. The model with an initially 175-km-thick lithosphere predicts

base of lithosphere

r 125 km 4.0 Adiabatic Boundary Condition • Ux= 5 mm/yr ' •'• • 120 0 Zo= 125 km

3.0 &•

crust/mantle

boundary

50

m.y.

1.0

.0 • 400 600 800 1000 1200 1400 1600

o

Temperoture C

Figure 6. P-T diagram for 125-km-thick lithosphere extending at 5 mm/yr, assuming an adiabatic lower thermal boundary condition. The model is the same as in Figure 4c. Much of the mantle initially lies in the supersolidus field for fluid-saturated

_ pyrolite, but little further melting is predicted because ascent trajectories do not move material further into the supersolidus field. The stippled pattern indicates the portion of the lithospheric mantle in which mafic components undergo partial melting during extension.

HARRY AND LEEMAN: PARTIAL MELTING OF SUBCONTINENTAL MANTLE 10,263

a)

Adiabatic Boundary Condition U x = S mm/yr Zo= 75 km

3.0

1.0

crust/mantle 60 boundary

40

60

80

100 m.y.

.o 4oo

b)

5.0

4.0

2.0

6OO

1.0

800 1000 1200

Temperature øC 1400 16oo

Constant Temperature Boundary Condition U x = 5 mm/yr Zo= 175 km

lOO

140

base of lithosphere

175 km

0

160 t

t

20

•4 40

crust/mantle

boundary

6O

.0 • • • 300 500 700 900 11 O0 1300 1500

o

Tempero•ure C

Figure 7. P-T diagrams for models with (a) 75-km-thick lithosphere and (b) 175-km-thick lithosphere. The models use the adiabatic lower thermal boundary condition. Other model parameters are those of models 2 and 3 in Table 2. The stippled pattern indicates the portion of the lithospheric mantle in which mafic components undergo partial melting during extension.

that no melting of mafic components in the lower lithospheric mantle occurs until after 20 m.y. (20% extension) (Figure 7b). The thickness of the region undergoing partial melting increases steadily thereafter, with increasing extension. This is similar to the pattern seen in the model with 125-km-thick lithosphere (Figure 6) but with the onset of melting being delayed in the thicker lithosphere.

The effect which the initial thickness of the lithosphere has on the timing of melt generation is summarized in Figure 8. Figure 8 includes the results of two models which are not shown here, which are similar to those in Figures 6 and 7 but with in- itial lithosphere thicknesses of 200 and 225 km. In the models in which the lithosphere is initially thicker than about 150-175 km, large amounts of extension are required before the onset of melting. Thick lithosphere is therefore unlikely to produce early syn-extensional magmas by melting of mafic components in the mantle. If the lithosphere is initially thinner than about 150 km partial melting occurs during the early stages of extension, with melting occurring from the onset of extension if the lithosphere is initially thinner than about 125 km.

The degree of melting is estimated by comparing the P-T models in Figures 6 and 7 with the empirical melting behavior for basalt reported by Thompson [1975], who also gives major element analyses of glasses in high pressure run products. The phase relations reported by Thompson are shown in Figure 9. Melt fraction is estimated for each run using measured concen- trations (Cl) of P205 and K20 in the glass phase and correspond- ing values (Co) for the bulk starting material. Assuming that P and K are partitioned strongly into the melt phase (i.e., bulk solid/melt partition coefficients, D, approach zero), then mass balance requires that under equilibrium melting conditions

Co/C t = El[1- O(1- F)]. (10)

where F is the melt fraction. For D=0, Co/C • = F. Because K and P are largely confined to the glass phase at all pressures, our estimates of F are insensitive to the crystalline phase assem- blages which vary from run to run. With the exception of a few runs, the independent estimates of F (Fp and FK) are in reason- able agreement. The variation in melt fraction with temperature above the solidus was assessed using a normalized temperature parameter, defined as T n = (T•- Ti)/(T • - Ts), where T•, T s, and T l

100

8O

4O

2O

0 100

i i I i •1 Ons'et '

.

i i i

125 150 175 200 225

Initial Lithosphere Thickness (km)

Figure 8. Variation in the timing of the onset of extension with initial lithospheric thickness. The model is that shown in Figures 6 and 7.

10,264 HARRY AND LEEMAN: PARTIAL MELTING OF SUBCONTINENTAL MANTLE

1500

1400

1300

1200

1100

OI+L

lOOO o 8 16 24 32

Pressure (kbar)

Figure 9. Phase diagram for an olivine tholeiite (59P13) from the Snake River Plain, Idaho [Thompson, 1975]. Abbreviations used are liquid (glass), L; olivine, O1; plagioclase, P1; clinopyroxene, Cpx; garnet, Ga; ilmenite, Ilm. Lowermost solid line approximately represents the initial melting conditions for this basalt.

are the respective experimental, solidus, and liquidus tempera- tures at the ambient run pressure. T• is well correlated with Fp but exhibits some scatter versus FI• (Figure 10); this scatter is largely eliminated by excluding data for four runs (see Figure 10 caption). Average melt fraction (Favg), based on the accepted run data, also varies systematically with T n. Linear fits to each of the data sets Fp and FI• (Figure 10) result in an average re- gression curve describing the melt fraction as a function of T,'

F•g = 0.217 + 0.680T,, (11)

with a correlation coefficient of 0.83. This relationship is ap- proximate at best, as it predicts nearly 20% melt at the solidus (T• = 0) and a maximum of only 90% melt at the liquidus (T,, = 1). In the absence of more quantitative high-pressure melting relations for mafic compositions, we have used this relation to represent the melting process. It provides a reasonable ap- proximation to fractional melting of a pseudo-eutectic material, for which a significant amount of melt will form near the solidus temperature. Near the liquidus, fractional melting of a eutectic material will leave an increasingly refractive residue, leading to a significant reduction in melt production (•F/•T).'

The total volume of melt can be estimated by integrating the melt fraction over the thickness of the portion of the lithosphere lying in the basalt supersolidus field at any given time:

z•(t)

Vm(t) = • Fv(z,t)dz. (12) z•(t)

Figure 10. Experimental melting relations for basalt [Thompson, 1975]. (a)-(c) Normalized run temperature T• versus estimated melt fraction F (calculated as described in the text) for (a) phosphorous and (b) potassium oxides and (c) the average of the two. Regression lines are based on the data indicated by the solid circles. (d) Open circles represent data which are inconsistent with the theoretical melting curve (see equation (11)). These points were excluded in the regression analyses. Vertical error bars indicate the standard deviation of the averaged data points. Error in T• is estimated to be _+5 %.

The rate of melt production is given by

Rm(t) = dV,, (13) dt

where z s' and z/are the depths of the basalt solidus and liquidus in the reference frame given in equation (2). Equation (12) is evaluated using Simpsoffs Rule, and the midpoint difference ap-

1.0 p

0.8

0.6

O.4 0.2 y = 0.240+0.655x R2= 0.83

0.0 ' ' ' ' ' ' ' ' '

F'/,' b' ' ' ' ' b) ß

O ß '

y = 0.163+0.738x R2= 0.84 ' 00 ' • ' • ' ' ' ' ' ' ß

ß ' ' ' ' ' Favg 0.8

0.6

0.4

0.2

0.0 0.0 0.2 0.4 0.6 0.8 1.0

0.8 0.6

0.4 0.2

10 ß ! ß I ß ! - ! - I

starting _%.^ composition

ß P205 equilibrium melting A K20

.1 I I I I • I ß I . i ß 0.0 0.2 0.4 0.6 0.8 1.0 1.2

Favg

HARRY AND LEEMAN: PARTIAL MELTING OF SUBCONTINENTAL MANTLE 10,265

proximation is used in equation (13) [Dahlquist and Bjorck, 1974]. It is worth noting that the melt production rates deter- mined from equation (13) neglect the effect which melt extrac- tion would have on the temperature of the model. Continuous removal of melt during the extensional process [e.g., Liu and Chase, 1991] could result in cooling of the lower lithosphere since some heat would be extracted with the melt. Equation (13) would then overestimate the amount of melt production. This effect would become more pronounced as extension pro- ceeds and progressively more melt is extracted. Melt production rates during the early stages of extension are less effected by temperature changes arising from melt extraction. In addition, if the base of the lithosphere is approximately isothermal (as ar- gued above and assumed in the models), then melt extraction would have no effect on the melt production rates. Neverthe- less, the model results should be viewed as representing maxi- mal melt production rates.

The rates of melt production predicted by the models in Fig- ures 6 and 7 are shown in Figure 11. The most obvious feature is the increase in the amount of extension required to produce melt as the initial thickness of the lithosphere increases. This was deduced previously from the melting relations in P-T space (Figures 6 and 7). In all of the models, the rate of melt produc- tion ultimately decreases (beginning immediately in the model with a starting lithospheric thickness of 75 km). The reduction in the rate at which magma is produced results partly from the fact that the rate of lithospheric thinning decreases with time (Figure 3b), reducing the rate at which mantle rocks move into the basalt supersolidus field. Another contributing factor is that the basalt solidus increases in slope at lower pressures (i.e., higher amounts of extension) and is nearly parallel to the ascent paths in the lowermost lithosphere (Figures 6 and 7). As a re- sult, the P-T trajectories of individual rock packets follow paths of constant normalized temperature (T n) at shallow depths and undergo little further partial melting. Furthermore, our calcula- tions ignore changes in bulk composition of the melt residue, which probably becomes more refractory with progressive melt- ing. A likely consequence is that melt production will diminish more rapidly than we estimate.

The model with the 75-km-thick lithosphere produces melt at the onset of extension, but the rate of melt production begins to decrease immediately (Figure 11). The model with a starting li-

150

100

50

i i i

•, I Lithosphere thickness I 75km

• I lZ3Km

0 20 40 60 80 100

Time (m. y.)

Figure 11. Melt production rate vs. time for models with initial lithosphere thicknesses of 75 km, 125 km, and 175 km. The extension rate is 5 mm/yr.

150

125

75

5O

25

0 75

, •.....• , , , 150 ,•/' ' ¾5'/. Vcak Magma

/• >'-'• Flux Ralc / 125

Time 1o Peak Magmati,•n •' -% '• 25 ! I I I • 0

1 O0 125 150 175 200 225

Initial Lithosphere Thickness (km)

Figure 12. Peak magma flux and the time until peak lithospheric magma production as a function of the initial thickness of the lithosphere. Extension rate is 5 mm/yr.

thospheric thickness of 125 km predicts a high rate of melt pro- duction from the outset and sustains high (but decreasing) rates of melt production for-20 m.y. Leemah and Harry [1993] showed that a similar model (using a constant temperature lower thermal boundary condition) was consistent with the early onset and duration of silicic mid-Tertiary magmatism and Miocene volcanism in the Great Basin. The model with 175-km-thick

lithosphere shows a delay in the onset of magmatism, and the maximum rate of melt production is lower than in the model with a 125-km-thick lithosphere. Optimal conditions for melt- ing of mafic material within the lower lithospheric mantle occur if the lithosphere is initially -100-150 km thick, with maximum melt production rates predicted for the model starting with a 125 km thick lithosphere (Figure 12). All models with lithosphere initially thicker than about 150 km require substantial amounts of extension before the onset of melting, and the resulting melts are produced at a lower rate than in the model with 125-km- thick lithosphere. All models with a lithosphere initially thinner than about 100 km produce melt at a relatively low rate, which decreases rapidly with continued extension.

Within the range of extension rates typical of most exten- sional tectonic provinces (between-2.5 and 10 mm/yr) the re- sults of the models are relatively insensitive to extension rate (see Figure 5). The onset and duration of magmatism (Figures 11 and 12) can thus be scaled to the amount of extension in or- der to extrapolate the results to faster or slower extension rates than used in the models. The magma flux rate scales similarly and is lower after a given amount of strain at lower rates of ex- tension. At extension rates lower than -2.5 mm/yr the adiabatic model may not be appropriate, and conductive cooling of the lower lithosphere may inhibit melt generation.

Discussion

The ascent of hot asthenospheric mantle has been proposed to account for melt production during extension in the Great Basin, either in consequence of steepening of the dip of the subducting Farallon plate, subduction of the East Pacific Rise, or the ap- pearance of a hot mantle plume [Snyder et al., 1976; Lipman, 1980; Thorkelson and Taylor, 1989]. As argued by Leemah and Harry [1993], this mechanism is unlikely to universally account for early synextensional magmatism because even unusually hot asthenospheric mantle will not melt unless the lithosphere is in- itially thinner than about 65-80 km [McKenzie and Bickle,

10,266 HARRY AND LEEMAN: PARTIAL MELTING OF SUBCONTINENTAL MANTLE

1988]. Furthermore, it is difficult to account for the evolved geochemical signature of early synextension magmas with as- thenospheric melts. Alternatively, the presence of volatiles within the lithospheric mantle may reduce the melting tempera- ture, promoting melting after relatively little extension (Figure 6). Such a melting scenario may account for early synexten- sional volcanism in many areas [e.g., Gallagher and Hawkesworth, 1992]. However, as discussed above the amount of melt produced in this way is limited because ascent trajecto- ries in the lithospheric mantle rocks are nearly parallel to the solidus of water saturated peridotitic rocks; as a consequence, the ascending mantle does not move further into the supersoli- dus field than it was prior to extension and little synextensional melting is predicted. Furthermore, volatiles are preferentially incorporated into the earliest melts. Unless they are continu- ously replenished, they will be rapidly depleted and the mantle will revert to dry conditions. For these reasons, we have exam- ined the potential for melting of mafic components within the li- thospheric mantle during extension.

In order to produce early synextensional magmas by partial melting of mafic components within the lithospheric mantle, ex- tension must involve portions of the mantle which are near the melting point of basaltic material (between 1100øC and 1333 ø C). These temperatures are predicted at-100-150 km depth if the thermal structure of the mechanically strong portion of the lithosphere is similar to the cooling plate model of Parsons and McKenzie [1978]. Such conditions are optimal for magma pro- duction by melting of mafic components in the lower lithosphere (Figure 12). Delamination of the lithospheric mantle [e.g., Bird, 1978; Houseman et al., 1981] may prevent preservation of geo- chemical heterogeneities in the lower lithosphere; however, if sufficient time passes following the delamination event (-50-75 m.y.; [Houseman et al., 1981]) the base of the remaining litho- sphere will heat up. The lithosphere may then resemble the thin (but warm) lithosphere used in model 2 (Figure 7a), and some melting of mafic components is predicted (Figure 12).

The inference that many Great Basin basaltic lavas are de- rived from lithospheric sources is supported by the fact that their chemical and isotopic compositions are distinctive from those of most MORB and OIB lavas and the interpretation by many that these characteristics cannot be explained simply as a result of assimilation of lithospheric components by ascending astheno- spheric-derived magmas [Ormerod et al., 1985; Fitton et al., 1988; Lure et al., 1989; Farmer et al., 1989; Menzies, 1989; Bradshaw et al., 1993]. There are exceptions to this view, for which either crustal or mantle wallrock interactions likely oc- curred [e.g., Glazner et al., 1991], but in detail these processes fail to explain the chemical and isotopic systematics of many ba- saltic suites [Lure et al., 1989; Kempton et al., 1991; Bradshaw et al., 1993]. For example, isotopic and trace element composi- tions of basaltic rocks in the eastern Great Basin transition zone

do not appear to lie on reasonable mixing trends involving de- pleted mantle and deep crustal compositions inferred from xenoliths and Proterozoic supracrustal rocks [Kempton et al., 1991]. Similarly, petrogenic models invoking crustal assimila- tion and fractional crystalization are unable to simultaneously match the elevated incompatible trace element compositions and Sr and Nd isotopic characteristics of basalts in the Snake River Plain and central Great Basin [Lum et al., 1989].

One consequence of the model proposed here is that the iso- topic and trace element compositions of melts produced from metasomatized lithospheric mantle should show spatial vari- ations correlating with major tectonic boundaries. In the Great

Basin, basalts from both the Idaho border region and from, southeastern California-southwestern Nevada display pro- nounced Sr and Nd isotope transitions coincident with the 87Sr/86Sr = 0.706 isopleth as defined by Mesozoic granitoids [Orrnerod et al., 1985; Leeman et al., 1992]. There is a prepon- derance of relatively high 87Sr/a6Sr and low 143Nd/144Nd ratios in primitive basalts erupted in eastern terranes, as compared to those erupted in the west. This correspondence presumably re- flects magma generation within the mantle straddling a signifi- cant pre-Tertiary tectonic boundary that marks the juxtaposition of Precambrian lithosphere to the east with accreted oceanic ter- ranes to the west. These respective isotope characteristics could only arise if the source distinctions have existed for times com- parable to the age of the lithospheric terranes, and their spatial distribution suggests that they reside within the North American tectonic plate. Interestingly, the basalts from the western Great Basin exhibit trace element compositions (e.g., elevated Ba/Zr) characteristic of many subduction-related lavas, despite the fact that they erupted millions of years after the cessation of subduc- tion at these latitudes [Ormerod eta. 1, 1985]. Apparently, the mantle sources were metasomatized by Mesozoic subduction- related fluids and then subjected to melting in the latest Ceno- zoic. This metasomatic overprint is seen in lavas from either side of the tectonic suture and is therefore not responsible for the isotopic heterogeneities inherited from the respective li- thospheric mantle domains. It is difficult to see how such het- erogeneities and metasomatic features could be inherited from a worldwide perisphere substrate over which the tectonic plate must move [e.g., Anderson, 1994]. In any case, if a long-lived geochemically distinct sublithospheric layer in the upper mantle is present, it should rise as the overlying mechanical lithosphere thins during extension; ascent trajectories in this layer would be similar to those presented above, resulting in similar melting relations.

Thus we view the lithosphere as an important source of ex- tensional magmas. For magma production to occur on the spa- tial and temporal scales observed in the Great Basin, and assum- ing that the model discussed in this paper is appropriate, it ap- pears that the preextension lithosphere was moderately thick (- 125 km), precluding significant melting of sublithospheric man- tle until the last few million years. We view the onset of mag- matism as fundamentally related to regional extension because such mechanisms as delamination or ascent of a plume require a substantial lag time between their initiation and the onset of magma production.

The largest variables in the models concern the amount of magma which is produced by partial melting of mafic material in the lower lithosphere and the relative timing between the pro- duction of melt predominantly in the lithosphere and the onset of magmatism arising from partial melting in the sublithospheric mantle. Factors controlling the volume of melt production from lithospheric sources which are not constrained by these models include the amount of mafic material available for melting within the lithospheric mantle, the effect which melt extraction has on reducing the temperature (and hence melt potential) during extension, the process of crustal anatexis and the amount of mafic magma required to produce the volume of crustal melts estimated to be emplaced in various extensional provinces, and the thickness of the mantle layer which is involved in exten- sional deformation. The models presented above have adopted simple assumptions which are best viewed as reasonable end- member conditions optimal for magma production arising from melting of mafic material within the lithosphere. Factors which

HARRY AND LEEMAN: PARTIAL MELTING OF SUBCONTINENTAL MANTLE 10,267

may contribute to an early onset of relatively voluminous melt- ing within the sublithospheric mantle that are not considered in this study include nonuniform extension (i.e., localized mantle upwelling), the role of small-scale convection in inducing en- hanced melt production after relatively minor amounts of exten- sion, and the possibility that passively upwelling sublithospheric mantle may initially lie at depths shallower than the 1333øC iso- therm. These factors may play a major role in contributing to the complex space and time patterns of magmatism in the Great Basin [e.g., Best and Christiansen, 1991; Axen et al., 1993].

Conclusions

The numerical models test the hypothesis that partial melting of mafic material entrained within melt metasomatized litho-

sphere may produce precursor magmas that result in crustal anatexis and the generation of large volumes of silicic to inter- mediate magmas during the early stages of extension. If the ba- sic tenents that the lower lithosphere contains a substantial amount of mafic material and that mantle as warm as 1300øC is

involved in extensional deformation are correct, then the models

confirm that the timing and geochemical characteristics of early synextensional magmatism can be explained as a consequence of melting of mafic components within the lower subcontinental mantle (SCM). The initial thickness of the lithosphere (taken to be equivalent to the conductive thermal boundary layer) plays a critical role in determining the amount and timing of magma generated by melting of mafic material within the SCM during extension. Little magma is generated if the lithosphere is in- itially less than -100 km thick because the ascent paths of lower SCM rocks are nearly parallel to the basalt solidus at pressures less than about 1.0 GPa. Conversely, if the lithosphere is thicker than -150 km, large amounts of extension are required to produce melt, and the rate of melt production is limited by the decrease in the rate of lithospheric thinning after large amounts of extension. An initial lithospheric thickness of-125 km is optimal for the generation of large volumes of early synexten- sional magma by melting of mafic components in the lower lithosphere. Under these conditions the deepest portion of the lithosphere lies near the basalt solidus prior to extension and crosses into the supersolidus field after minor amounts of ex- tension. Melting is limited to the lower 25 km of the SCM be- cause shallower portions of the lithosphere are at temperatures too low to ever cross the basalt solidus. Magmatism arising from such a melting scenario is transient and may last up to 30- 40 m.y. The duration of the magmatic episode and the amount of extension prior to the onset of melting increase with greater initial thicknesses of the lithosphere. The space and time pat- terns of silicic to intermediate synextensional magmatism and the trace element and isotopic characteristics of Miocene basalts in the Great Basin are consistent with partial melting of melt metasomatised lower SCM if the preextension thickness of the lithosphere was -125 km thick and the base of the lithosphere was warmer than - 1100-1300øC.

Appendix

The equation governing heat transport in individual rock packages is developed in a Lagrangian coordinate system [z', t}, where z' indicates the particle of rock. The heat energy in an elemental volume of rock is

Q=pCT(z',t) (A1)

and the rate at which heat energy in the elemental volume changes is

- 3Q 3T(z',t) (A2) q =-•-= pC 3'••" The conductive heat flux through the elemental volume is

•c = K i)T(z' ,t) (A3) 3z'

and the rate of change of heat energy arising from heat flux through the elemental volume is

•)2r(z' t) V ß •c = g•' (A4) •z, 2 ß

The amount of heat energy produced by adiabatic pressure changes is

Qa = PCfiTa (A5)

where fit a is the adiabatic temperature change. It is convenient to define the adiabatic temperature change in terms of the Eule- rian coordinate z, which describes the position of the elemental rock volume relative to the surface of the Earth. If the element

ascends a distance 3z, then the adiabatic heat energy in the {z, t } reference frame is

Qa = [3C3z 3Ta (z,t) 3z = [3C3zG a (A6) where G a is the mantle adiabatic temperature gradient. The rate of adiabatic heat energy change within the elemental volume is

•Qa •z 3t =oC• 'Oa' (A7)

If the thickness of the lithosphere at any time is given by h(t), then the Lagrangian and Eulerian coordinate systems are related,

z = z' h(t), (A8)

and the rate of ascent of the rock packet is

3z 3h(t) (A9)

Combing equations (A2), (A4), (A7), and (A9) and including a radioactive heat generation rate A(z') within the element leads to the Lagrangian form of the equation governing heat balance within each moving particle of rock:

c 3T K 32T p •--•-z ,= A(z' )+ +OCUzG a . 3z '2 (A10)

Acknowledgments. We thank E.I. Smith, M. Liu, and D.L. Anderson for their constructive reviews.

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D.L. Harry, Department of Geology, University of Alabama, Box 870338, Tuscaloosa, AL 35487-0338. (e-mail: dharry@geophys .geo.ua.edu).

W.P. Leeman, Department of Geology and Geophysics, Rice University, P.O. Box 1892, Houston, TX 77251. (e-mail: leeman @ g eoph ysic s.rice.edu)

(Received May 13, 1994; revised October 31, 1994; accepted November 21, 1994.)