Thermal and kinematic model of the southern Rio Grande rift: inferences from crustal and mantle xenoliths from Kilbourne Hole, New Mexico

373

Tectonophysics, 197 (1991) 373-389 Elsevier Science Publishers B.V., Amsterdam

Thermal and kinematic model of the southern Rio Grande rift: inferences from crustal and mantle xenoliths from Kilboume Hole, New Mexico Gilles Y.A. Bussod ’ and David R. Williams Department

’

of Earth and Space Sciences, University of California, Los Angeles, CA 90024, USA

(Received March 9.1990; revised version accepted May 14.1990)

ABSTRACT Bussod, G.Y.A. and Williams, D.R., 1991. Thermal and kinematic model of the southern Rio Grande rift: inferences from crustal and mantle xenoliths from Kilboume Hole, New Mexico. In: A.F. Gangi (Editor), World Rift Systems. Tectonophysics,

197: 373-389.

A simple thermal conduction model constrained by piezothermometric data obtained from coexisting minerals in both crustal and mantle xenoliths from Kilboume Hole maar, New Mexico, is used to understand the development of rifting and the thermal evolution of the lithosphere in the southern Rio Grande rift. The one-dimensional finite difference code models crustal thinning and underplating by magmatic injection into the lower crust and uppermost mantle. The model is in good agreement with crustal and mantle thermometric arrays assuming a transient, pm-rift gradient superimposed on a stable geotherm and yielding a net gradient of 20°C km-’ in the 40-km-thick crust and 3’C km-’ in the upper mantle 30 million years ago. This is consistent with pm-rift volcanic activity in the area possibly related to Laramide erogenic events. A massive injection of basaltic dikes and sills (= 70 vol.%) in the uppermost mantle and at the base of the crust probably occurred in the early development of the southern Rio Grande rift, between 20 and 30 Ma. The heat contribution due to the injection of magma is treated as a temperaturedependent heat source distributed uniformly throughout the intruded region which extends from 25 to 45 km depth. Textural observations along with major and trace element compositions of the mantle and crustal xenoliths indicate that metasomatism and recrystallization may have occurred during liquid infiltration and partial melting 20 to 30 Ma. This is consistent with early rift bimodal volcanism associated with crustal contamination and linked to approximately 25% total NE-SW extension. Die injection and CNStd thinning ceased 20 Ma associated with the mid-Miocene lull in volcanic activity. The total volume of material injected, 100 kms per km length of rift per million years, is approximately an order of magnitude less than for a typical mid-ocean ridge. Conductive cooling of an uppermost mantle (40-45 km depth) from temperatures of over 1200°C (22 Ma) leads to the formation of a granulite metamorphic crustal complex and generation of a surface heat flow of 82 mW m-*, in accordance with observed present-day values in the southern Rio Grande rift. Overall, the results from the dike injection model are consistent with both the crustal and mantle piexothermometric arrays and present-day geochemical and geophysical data from the southern Rio Grande rift.

Intmduction The southern Rio Grande rift (SRGR) is located between two main physiographic and crustal provinces; the Colorado Plateau and southern Basin

’ Present

address: Bayerisches Geoinstitut, Universitgt Bayreuth. Postfach 10 12 51. 8580 Bayreuth, Germany. 2 Presentadd-: Carnegie Institution of Washington, DTM, 5241 Broad Branch Rd., Washington, D.C. 20015, U.S.A. 0040-1951/91/$03.50

and Range to the west, modified by late Cenozoic tectonism, and the unmodified North American craton to the east. The SRGR merges with the Basin and Range in southern New Mexico and extends southeast into west Texas and Chihuahua, Mexico (Seager and Morgan, 1979; Olsen et al., 1987). Kilboume Hole is one of several late Pleistocene basaltic maars found in the central part of the Potrillo volcanic field and contains a large variety of crustal and mantle xenoliths associated with basanitic lavas dated at 80 (f 10) Ka by apatite fission track (V. Harder, pers. commun.,

8 1991 - Elsevier Science Publishers B.V. All rights reserved

374

1989). The maar is located of the SRGR north-south road”

nearly “axial”

refraction

and constitutes

along

the central

at the intersection and the east-west

profiles a natural

axis

of the “pipeline

tion based and

gradient

from Sinno et al. (1986),

spheric

structures

bore-hole

an extended

servations

sections,

on the Kilbourne

on current

geophysical

is coupled

into the mod-

P-T-time

evolution

informa-

Hole xenolith and petrological

with a comprehensive

mal model in order to understand tectonic

et al., 1986; Morgan

perature

ern lower crust and upper mantle. In the following

1984; Aldrich

lead us to conclude

the thermal

of the continental

suite obtherand

lithosphere

were substantially

period

A number

of pre-rift

of sources

a 10 m.y. rhyolite

tem-

and litho-

modified

magmatism.

in our thermal point

Ma as representing

stage extension

stable

as well as the crustal

modification is included an initial condition. 27-32

et al.. 1986).

that the original

by This

model

as

to approximately

the beginning

in the Rio Grande

of early

rift followed

period

of massive

bimodal

volcanism

(e.g.,

Chamberlin,

by

basaltic1978;

beneath the SRGR. The 1-D finite difference thermal model incorporates crustal thinning and dike

Chapin, 1978, 1979; Eaton, 1979; Seager et al., 1984; Aldrich et al., 1986; Morgan et al., 1986;

injection, and is constrained at depth by thermobarometric data obtained from lower crustal and

Olsen et al., 1987). This stage of rifting

upper mantle

xenoliths.

comprises

nearly 80% of all southern rift volcanics and is intimately associated with low-angle extensional faulting (Morgan, 1986). In our thermal model we

Tectonic setting

assume

The base of the crust in the SRGR consists of intraoceanic island arc and marginal basin litholo-

jection at 30 Ma, and a 10 m.y. duration stage rifting in the southern Rio Grande ding 20 Ma. Most details of our results

gies accreted to the North American continental margin between 1.8 and 1.6 Ga (Condie, 1982). Subsequent periods of talc-alkaline arc magmatism and eastward migration of volcanism from the continental margin 75 to 60 Ma are associated with subduction/accretion events (Lipman et al., 1986). Increased magmatism between 75 and 50 Ma appears to be related to an abrupt change in plate convergence rates (Engebretson et al., 1984) and a possible flattening of the subducted plate (Coney and Reynolds, 1977). Middle Tertiary volcanism in the Rio Grande rift and in the Basin and Range to the west corresponds to a period of transition from convergent margin/continentalinterior arc magmatism to extensional continental tectonics, and may have resulted from the collision of the East Pacific Rise with the North American plate approximately 30 Ma (Atwater, 1970). We assume that the region prior to rifting ( = 50 Ma) was analogous to the present-day Great Plains (see Cook et al., 1979; Morgan et al., 1986). From this time until about 30 Ma, this “Great Plains” structure was modified by magmatism and crustal thinning (Elston and Bornhorst, 1979; Olsen et al., 1987). Intense magmatic activity in the area, ranging from 45 to 30 Ma (Seager et al.,

sensitive

the initiation

of rifting

to the time chosen

and magmatic

in-

of early rift, enare not

for the initiation

of

rifting and magmatism. The initial stage of extensional tectonics was followed by a period of quiescence, or a “midMiocene lull”, of 10 million years duration, possibly related to the early development of a transform boundary along the west margin of North America (proto-San Andreas fault) and extreme extension in the Basin and Range. We infer that after 20 Ma, only stable mantle heat flow, radiogenie heating, and conduction occurred until the present time. This may be an oversimplification as late stage upper the mid-Miocene

crustal magmatic events followed lull (Chapin, 1979; Eaton, 1979).

However, the effect of these events on the development of the southern rift is assumed to be minor. Limited extension characterized by highangle normal faulting and relatively minor lower Miocene to Holocene alkalic basaltic volcanism distinguishes the SRGR from the Central and Northern rift provinces, where tholeiitic basalts erupted from central volcanic complexes from 13 Ma to the present (Baldridge et al., 1989). However, the SRGR may be more representative of the early rift history as the tholeiites from the central rift are generally associated with an anomalous

THERMAL

AND

KINEMATIC

MODEL

OF THE

SOUTHERN

RIO GRANDE

north-northeast trending linear volcanic array Jemez Lineament, located at the boundary of Colorado Plateau core and a transition zone to southeast (Baldridge et al., 1989; Aldrich et 1986).

375

RIFT

TEMPERATURE (‘C)

or the the al.,

,‘I”

q

,yl

,“T”

,8tfO ,ly,,,,l4/

L

Vp (Km/s)

I PLAGIOCLASE

20 -

Thermobarometric

estimates

Details of the thermobarometric determinations on the Kilboume Hole xenolith suite are presented elsewhere (Bussod and Irving, 1990). Estimation of equilibration temperatures and pressures recorded by mineral assemblages in xenoliths is fraught with difficulty because of inaccuracies in their calibration, differing P-T blocking conditions for different mineral reactions, and matrix effects in their application. Nevertheless, based on the careful evaluation and application of available thermobarometric techniques suitable to the xenolith assemblages from Kilboume Hole, determination of the important constraints on the thermal history and composition of the lithosphere beneath this area is possible. Shown in Fig. 1 are the piezothermometric arrays determined from the crustal granulites and mantle spine1 lherzolites from Kilboume Hole. This data along with the occurrence of upper crustal assemblages within the maar indicate that the lithosphere has been sampled continuously from a depth of 70 km. The xenolith suite from Kilboume Hole is therefore assumed to be a representative sample of the lithologies which characterize the central SRGR.

z

40 -

s

RE%E PRECISION

E 4 o

60-

80 -

0

CRUSTAL GRANULITES LHERZOLITES

0

FINE

0

COARSE COMPOSITES

-

OUVINE - SPINEL TEYPERATLIRE RANGE

(99)

100 x (h@Mg+Fe) 1

Fig. 1. Kilboume

I

\ *=\ w

@E

100 -

\

I

I

OLMNE

I

I

I

granuloblastic

i

I

I

I

I

granulites (circles)

data. P-T

and protogramtlar

mantle lherxolites

Composite

represent

olivine

lherxolite xenoliths

bols). Arrows

connected

the estimated

sample

olivine-spine1

using

and

are from Padovani

Mg numbers

Numbers

in lhetzolites.

are also shown (half open sym-

to thermobarometric

tines represent

estimates

(squares)

(1977) and Bussod and Irving (in prep.), respectively. in parentheses

I

Hole lower crustal (solid circles) and upper

mantle (open symbols) piezothermometric for the crustal

I ]

data points by

range in temperatures thermometry

for each

(Fabribs,

1979).

Curves I and 2 delineate the stability field of spine1 lherzolite. Crustal compressional

wave velocities

et al. (1986). Sample numbers xenoliths (open circles). T-P

are derived from Sinno

refer to granuloblastic relative precision

mantle

of the therm*

barometric estimates for all samples is also shown.

The crustal suite

The lower crustal xenoliths from Kilboume Hole are mafic and include garnet-bearing paragneisses, two-pyroxene-bearing orthogneisses, charnockite, and anorthosite (Padovani and Carter, 1977). The crustal piezothermometric array is based on garnet granulites consisting of strongly foliated rocks (granuloblastites) with almandinepyrope garnet, sanidine, quartz and silhmanite (Padovani and Carter, 1977). These assemblages are believed to represent a Proterozoic metamorphosed pelitic crust (Padovani and Carter, 1977), subsequently reheated by a later thermal

event related to rifting (Bussod and Irving, 1981; Padovani and Reid, 1989). Piezothermometric estimates are from Padovani and Carter (1977), and are based on Ca-exchange equilibria between garnet and plagioclase (Ghent, 1976), and on K, Na, and Ca equilibria between potassium feldspar and plagioclase (Stormer, 1975). More recent thermobarometric calibrations applied to these samples are in good agreement with these estimates (Bohlen et al., 1983). Ahhough P-T uncertainties are difficult to ascertain, the relative precision for the piezothermometric data is approximately f 0.1

376

GPa pressure and f50°C, based on mineral assemblage (Padovani, 1977). The crustal velocity structure (Fig. 1) is also consistent with the P-T error estimates (Keller et al., 1979; Olsen et al., 1979; Sinno et al., 1986). The crustal thermobarometric data form an array (Fig. 1). which has been interpreted to represent present-day in-situ conditions (Padovani and Carter, 1977; Olsen et al., 1979; Padovani and Reid, 1989). This implies an average present-day geothermal gradient of 30-35°C km-’ for the SRGR and a partially molten crust-mantle boundary. This interpretation is principally based on the homogeneity of coexisting mineral compositions, a high present-day average regional heat flow (100 mW mv2) and anomalous P,, upper mantle seismic velocities of 7.6-7.8 km s-’ (Olsen et al., 1979). However, several observations suggest that the crustal array represents peak metamorphic- and not present-day conditions: (1) The anomalous MOHO velocity in the SRGR is not correlated with high seismic attenuation (Davis, 1987), commonly associated with the presence of a liquid phase. (2) The present-day heat flow along the rift axis and in the vicinity of Kilbourne Hole (80 + 10 mW me2) is lower than the average regional value which may include hydrothermal circulation in the rift flanks (Reiter et al., 1986). (3) Garnet terrains throu~out the world exhibit similar average geothermal gradients (dT/dZ ~22%35°C km-‘), regardless of their emplacement history. This is p~ncip~ly due to the low cation diffusivities and reaction kinetics in Iithologies of this composition (Elphick et al., 1985). The garnets from these assemblages commonly exhibit broad, homogeneous core compositions and thin zoned rims (Bohlen, 1987). However, garnet rims from Kilboume Hole granulites have been replaced by glass and quench crystals formed by decompression melting (Padovani, 1977) such that their cooling history has been obliterated. Conversely, exsolution features are common in feldspar and pyroxene mineral “cores”, indicating that higher temperatures and/or possibly higher pressures predated the “rim” equilibration conditions (Padovani and Carter, 1977). Although local equilibrium can be maintained dur-

G.Y.A. BUSSOD AND

D.K. WIl.l.IAMS

ing initial retrograde thermal events (Bohlen, 1987), for relatively rapid cooling rates crustal granulite assemblages do not necessarily record the modem cooling history as slow cation diffusivities and reaction kinetics do not allow mineral equilibration (Elphick et al., 1985; Bohlen, 1987). Based on these observations, the crustal geothermometric array is interpreted as representing relatively recent (d 20 Ma) “peak” metamorphic thermal conditions, and not present-day conditions. The mantle suite The upper mantle xenoliths from Kilbourne Hole are dominantly spine1 lherzolites. Although commonly regarded as “typical” upper mantle samples, they span a wide range of mineral modes, textures, isotopic-, major- and trace-element compositions (Irving, 1980; BVSP, 1981; Roden et al., 1988; Bussod and Irving in prep.). The spine1 lherzolites used in the determination of the mantle piezothermometric array are divided into two textural types (Harte, 1977): granuloblastic tabular (fine grained) and protogranular (coarse). Both are occasionally cut by spine1 pyroxenite veins (in composite xenoliths), interpreted as the products of polybaric fractionation from basaltic intrusives at depth (Irving, 1980). Equilibration temperatures estimated for the mantle samples are based on two independent calibrations using independent phase assemblages: the temperature-dependence of the pyroxene miscibility gap (Wells, 1977), and the partitioning of Mg-Fe between coexisting olivine and spine1 pairs (Fabrib, 1979). Equilibration pressures were estimated using the two-pyroxene thermometer {Wells, 1977) in conjunction with a thermobarometer based on the solubility of Ca in olivine coexisting with ortho- and clino-pyroxene (Finnerty and Boyd, 1978; FiMerty and Rigden, 1981; Adams and Bishop, 1982, 1986; Finnerty, 1989; Kiihler and Brey, 1990) that is empirically calibrated to the natural system (Bussod and Irving, in prep.). The P-T estimates (Fig. 1) generally agree with previous determinations for the same locality (Reid, 1976; Mercier, 1977, 1980), and based on an overview of synthetic and natural systems are within the stability field limits for

THERMAL

AND

KINEMATIC

spine1 lherzolites

MODEL

at f 0.15 GPa and

determinations

f 30°C (Bussod

prep.). Based on thermobarometric two-pyroxene 1990),

(Bussod

the granuloblastic are

spatially

shallower

depths

(34-54

xenoliths

and

more, the granuloblastic

and Irving,

using the

and and

in

samples

similar

protogranular

and lherzolites disThe

generally the more xenoliths in prep.). also show their con-

described by Mercier (1977, 1980) for the adjoining Basin and Range and Colorado Plateau provinces, which suggests some lateral continuity to (Reid, 1976;

Mercier, 1977) the mantle array (within error) predicts temperatures lOO-200°C lower than the crustal array at a depth of 30 km (Fig. 1). However, this apparent discrepancy is resolved when considering the thermal history recorded in the mantle xenolith assemblages. The temperatures derived from two-pyroxene and olivine-spine1 mineral

pairs (Fig.

cores of spine1 grains

1) are in good agreement

for

the protogranular xenoliths (AT = 50°C) and yield apparent equilibration temperatures between 950°C and 1060°C (Bussod and Irving, in prep.), whereas large temperature discrepancies are found using the two methods in the granuloblastic and composite rocks. The temperature estimates using olivine-spine1 pairs in the composites and three of the four granuloblastic rocks, are systematically higher than temperature estimates using coexisting pyroxene pairs. In addition, the oxide phase is demonstrably out of equilibrium with the silicate phases for these samples (Bussod and Irving, 1981). However, neither this disequilibrium nor the inherent differences between the thermometers can

temperature zoning

range

in spine1

sent in all xenoliths,

grain

sizes.

and transport

reaction

Lobate are pre-

the zoning

perturbations

liths by the host basanite. and silicate mineral

even though both samples show a

indicating

with the incorporation

in the

to the granu-

with local melting

due to late stage thermal

cate different

differences

patterns

are restricted

loblastic and composite xenoliths protogranular and granuloblastic

mantle are from

counterparts.

investigators

large

comparable

stituent phases, represented by the olivine compositions in Fig. 1. Similar relationships have been

the west. As noted by previous

such

spine1 edges associated

are compositionally

former are compositionally fertile and have higher bulk Fe and LREE than barren MORB-related protogranular (Irving, 1980; Bussod and Irving, Granuloblastic and composite samples a wider range in the Mg/Fe ratios of

for

Irving,

km) (Fig. 1). Further-

lherzolites

from composite to their

account

(AT = 3OO’C). Retrograde

km) than the coarse pro-

(50-67

371

RIFT

The relative

composite

associated

RIO GRANDE

is estimated

estimates

thermometer

xenoliths togranular

SOUTHERN

of this composition.

of the P-T

precision

OF THE

is not

associated of the xeno-

These observations

kinetics

between

indi-

the oxide

phases and are consistent

with

an earlier thermal event affecting only the granuloblastic and composite rocks. This thermal perturbation reconciles the crustal and mantle piezothermometric arrays (Fig. 1) and indicates that an inverted

thermal

gradient

once existed

in the up-

per mantle beneath Kilbourne Hole. Consistent with mineral homogenization times (Bussod and Irving, in prep.), the inverted gradient is assumed to be associated with a relatively recent

thermal

restricted

to

perturbation the

(< 25 Ma), spatially

crust-mantle

boundary,

and

therefore is inconsistent with a simple model of lithospheric thinning by asthenospheric upwelling. Assuming

that the thermal

event

recorded

in the

xenolith mineral assemblages from Kilbourne Hole is the result of a regional metamorphic process associated with rifting, we have developed a numerical model which incorporates crustal thinning and injection of magma into the lower crust and upper mantle of the SRGR. The concept large intrusive complex in the lower crust

of a and

upper mantle of the SRGR is further supported by the unusually widespread occurrence of intrusive xenolith lithologies (composite xenoliths and mafic megacrysts) this area (Padovani Baldridge, Thermal

found in many localities and Carter, 1977; Fodor,

1979; Irving,

1980; Dromgoole,

from 1978; 1984).

model

The thermal evolution of the center of the southern Rio Grande rift is calculated using the following equation:

378

G.Y.A.

where T is temperature. f is time, z is depth, K is thermal diffusivity (= k/pC,, where k =

estimated difference

conductivity.

lateral

and

right-hand advection heat

p = density,

o is advective

side represent

certain injection

and

heat

terms

respectively sources

contexts

hereafter

of magma

effects.

a one-dimensional

ference

scheme.

The

scheme

is equivalent

magma although

by

of material

an

spreading

crustal

and

solved on an Eulerian

is solved dif-

the

finite

through

thinning.

difference

through the the material. the grid due to

Equation

grid. The grid spacing

and time steps (At)

were chosen

effect of third-order

terms in the conduction

(1) is ( AZ)

to minimize

the solu-

tion. The values used in this calculation are AZ = 1 km and Ar = 5 ka. The value K is 1.11 X 10 --’ m”

that the amount

in the crust and 1.06 x 10ph m* s- ’ in the mantie. The advective term in eqn. (1) represents the upward movement of material due to spreading concurrent with dike injection and the attendant S

symmetric about its axis. Calculating the characteristic time, T = d2/tc, for conduction over a d for geologic materials, shows this to be assumption. We have also numerically

km

is transported

and movement

of heat conducted laterally out of the region of dike injection is small, and that the rift itself is

distance a good

lO”C, which

to these calculations.

in

use of a one-dimensional to assuming

loss is approximately

scheme in three ways: conduction material, injection of dikes within

form will have forward

thermal

Heat

AN13 D K W,,_I.,AMS

the effect using a two-dimensional finite grid and found that the maximum

is not significant

to magma

The equation explicit

the

radioactive

due

as dikes,

in any other

on

conduction,

thinning,

We will refer to the intruded

the same thermal using

The

due to lithospheric

sources,

injection.

and Cr = heat capacity),

velocity.

HUSSOD

’

thinning

of the crust.

It is assumed

that all thin-

TEIIP

a

KH

0

mantle

j

dike Ton

t----22

km

j

: 800-c km---i

33

1050T

60

Fig. 2. Finite difference thermal model. (a) Initial conditions 30 Ma: no extension or intrusive% and the crust-mantle boundary is at 40 km depth. The box represents the initial intrusive area. (b) Final configuration after 10 m.y. inlrusive history (20 Ma): 25% extension

and 70% intrusives.

The original crust-mantle

boundary

(banded box) is at 26 km depth. Kilbourne Hole location (KH),

is at 31 km depth and the top horizon model densities and approximate

of underplated

temperatures

region

are also shown.

THERMAL

AND

KINEMATIC

MODEL

OF THE

SOUTHERN

RIO GRANDE

ning takes place above the dike injection region, and that upward movement of the dike injection region itself as well as the area below the dikes is simply due to this thinning (Fig. 2). The effect of advection on the temperature distribution is evaluated using an upstream differencing method. We have limited the total volume of dikes injected per km of distance along the rift to 1000 km3, evenly distributed within a predefined dike injection region. The heat due to dike injection, H, is represented by a temperature-dependent heat source:

HZ+ P

[

L+

(Th.4

-

TLWPL

RIFT

379

heating due to radioactive decay in the crust and mantle. This term is represented in the model as a stable temperature distribution, upon which the other temperature contributions are superimposed. A plausible heat production distribution is used to calculate an equilibrium temperature profile, which is assumed to be maintained throughout the history of the rift. The model also includes the effect of pre-rift heating. This is represented as a transient temperature profile superimposed on the stable profile before the actual initiation of dike injection and rifting. This pre-rift temperature profile is simply added to the time-dependent portion of the temperature as an initial condition. Parameters

+(r,-T,)(

ci$L)+(=s-T)Cp]

(2)

where y is the proportion of material injected per unit time step, C, and CpL are the solid and liquid heat capacities, respectively, L is the latent heat of fusion, and TM, TL, and T, are the injection, liquidus, and solidus temperatures respectively. The total heat input from the dikes is a function of the ratio of mass of injected material to mass of surrounding rock, the temperature at which the dikes are injected, the temperature of the adjacent country rock, and the thermal parameters representing the various materials involved. The dikes cool from their initial injection temperature, releasing heat to their surroundings. Between the liquidus and solidus temperatures, latent heat of crystallization is also released as the dikes solidify. The material continues to cool, releasing heat to its surrounding until the temperatures equilibrate. The dikes are assumed to be emplaced rapidly, with negligible loss of heat from the material to the mantle until the predefined dike injection region is reached and the magma halts its ascent. Therefore, the dike injection term in the equation is only applied within the dike region. Note, however, that spreading and crustal thinning will cause this region to move upward with time. The heat from the dike injection is assumed to be distributed over the entire dike injection region, and the dikes are assumed to reach thermal equilibrium with the country rock in one time step. The remaining term (A) in eqn. (1) represents

In order to apply our thermal model to the evolution of the SRGR, we must determine or estimate values for the relevant parameters (Table 1). In the following discussion, we will attempt to justify the values chosen, and discuss the effects of changing these parameters on the final results. The timing of events used in our model is principally based on the tectonic history of the region. The geologic structure is modelled at 30 Ma (Fig. 2a) and at 20 Ma (Fig. 2b), before and after magmatic injection, respectively. Our thermal model has its beginning at 30 Ma, the initiation of rifting. At this time, the crust is assumed to be 40 km thick, consisting of a 30 km upper crust, and a 10 km thick lower crust. This value is intermediate to that for the present-day Colorado Plateau to the west and the Great Plains to the east, initially approximately 50 km thick (Stewart and Pakiser, 1962; Keller et al., 1979; Cook et al., 1979), and thinned by about 10 km due to pre-rift activity. Our lowest crustal layer is equated with the southern Great Plains layer determined by Keller et al. (1979) using surface wave dispersion, which after crustal thinning up to 30 Ma occupies the depth range from 30 to 40 km. The density for the upper crust is taken to be 2700 kg rnT3 (cf. Keller et al., 1979; Cordell, 1982; Daggett et al., 1986). Measured densities of lower crustal granulites from Kilboume Hole range from 2900 to 3100 kg rnp3 (Padovani and Reid, 1989). Keller et al. (1979) using surface wave dispersion estimate a

3RO

Ci Y A BUSSOD AND

I, K WILLIAMS

TABLE 1

value of 2900 kg m-3

Parameter values used in the thermal model

Great Plains, and we take it to represent the lower 10 km of our modified crust. We assume a mantle

Model parameter K/t

Value

References

history

Inception of rifting Cessation of rifting Total extension Percentage dikes initial

30 Ma 20 Ma 50 km 70%

a, e, f, 1. u. z a, e, f. 1. u, 2 u

40km 30 km 10 km

g. 4

studies

initial

Reiter

“C-’

(e.g., Decker

et al., 1979;

1988), although

the values

cluster

3.0 W m-

in the most recent

about

a mean

of ap-

’ “C -‘. which we choose as

S

b (p.1153)

8x10-.” W kg -’ 10-i’ W kg-’ IO-” W kg-’ 10 km 32 mW me2

c c p c c,j, t

r, v. x

u. x

u, x

value for the thermal

conductivity

aa, k, m, n aa. n, 0

of

upper crustal and mantle values. This allows us to use a constant thermal diffusivity for the whole crust. For the mantle, a slightly higher thermal conductivity of 3.5 W m-’ OC-’ is assumed, whereas assumed

a heat capacity for all (solid)

of 1000 J kg-’ “C- ’ is material in this model

(BVSP, 1981, pp. 1153 and 1135, respectively). We assume that the initial stable geotherm remains constant throughout the 30 m.y. represented by the model. We have assumed a surface heat production of 8 X lo-” W kgg’ with a characteristic depth of 10 km, in accordance with Blackwell (1971). Decker and Smithson

(1975) and

Cook et al. (1978). A constant heat production of lo-” W kg-’ is used throughout the lower crust

1450°c 13cQ”C 1200°C 8x10s J kg-’

0

(Blackwell,

d

1980). The upper mantle heat production is taken to be constant to 120 km with a value of lo-” W

IGO0J kg-’ “C-’ 1.500J kg-’ “C-’

b (p.1135) y

kg-’ (Kaula, 1968). A boundary condition

Heat capacities

Solid Liquid

Kelley

1986; Reiter et al., 1986; Nathenson

the upper crust in our model. For convenience, we assume a conductivity of approximately 3.2 W m - ’ “C ~ 1 for the lower crust. part way between

Dike material (dry hasanite)

Injection temperature Liquidus (1.5 GPa) Solidus (1.5 GPa) Latent heat

conductivi-

a wide range of

a reasonable

3.0 W m-’ “C-’ 3.2 W m-’ Y-’ 3.5 W m-’ “C-l

20°C km-’ 3°C km- ’

1975;

thermal

rift exhibit

et

(Keller

h, i, q

transient gradients

Crust Mantle

crustal

of the

9. w h. q

Heat produciion

Surface Lower crust Mantle Characteristic depth Mantle heat flow

and Duncan,

proximately 27OOkgm-’ 2900 kg m-s 3300 kgm-’ 3000kgme3

of upper

Smithson,

and Guffanti.

Thermal conductiuities

Upper crust Lower crust Mantle

Estimates

ties for the Rio Grande

values from 2.8 to 3.3 W m-’

DPP?Si
Upper crust Lower crust Mantle injected magma

density of 3300 kg m- ’ for our model al., 1979; Cordell, 1982).

and

thickness

Crustal thickness upper crust Lower crust

for the lower crust

References: a = Aldrich et al., 1986; aa = Arculus, 1975: b = Basaltic Volcanism Study Project, 1981; c = Blackwell, 1971: d = Bottinga and Allegre, 1978; e = Chamberlin, 1978; f = Chapin, 1978; g = Cook et al., 1979; h = Cordell, 1982; i = Daggett et al., 1986; j = Decker and Smithson, 1975; k = DePaolo, 1979; I= Baton, 1979; m = Gbiorso and Carmichael, 1980; n = Green, 1973; o= Irving, 1974; p = Kaula, 1968; q = Keller et al.. 1979: r = Kelley and Duncan, 1986: s = Kushiro, 1980; t = Lachenbruch and Sass, 1978; u = Morgan et al., 1986; v = Nathenson and G&anti, 1988; w = Padovani and Reid, 1989; x = Reiter et al., 1986; y = Richet and Bottinga, 1986: z = Seager et al., 1984.

1971; Cook et al., 1978; Sclater

for heat

et al..

flow into

the

lithosphere from the mantle at depth must also be assumed. Blackwell (1971) uses a value of 30.5 mW m-* for his stable continental model which Cook et al. (1978) apply to a Great Plains geoDecker and Smithson (1975) and therm. Lachenbruch and Sass (1978) each estimate a reduced heat flux for the southern Rio Grande rift of 33.4 mW rnm2. For our thermal model, we take a value of 32 mW m-’ for the heat flux coming into the base of the thermal lithosphere at 120 km. Because the heat is being efficiently removed to

THERMAL

AND

KINEMATIC

MODEL

OF THE SOUTHERN

RIO GRANDE

Tempmtufe("C) 0

200

400

600

600

1000

1200

1400

'I"'I"'I"'I"""'I"'

381

RIFT

greater than 3°C km- ‘. We have assumed transient temperatures superimposed on the stable geotherm

yielding

gradients km-’ this

more

of 20°C

conservative

km-’

in the uppermost exact

transient

mantle profile

match

the piezothermometric

dicate

that

required.

a pre-rift

of material

of the crust and upper initial

Great

gradients

,oo’,,,‘,,.‘...‘...‘...‘~.. Fig. 3. Assumed stable and transient southern Rio Grande

the crust-mantle flow

remains

’

model geotherms for the

rift at the start of the model (30 Ma).

boundary, constant

we assume

throughout

the

this heat 30 m.y.

history of the rift. The upper boundary condition is a constant temperature of 0°C at the surface. The values chosen here give us a stable geotherm (Fig. 3) on which changes due to dike injection and crustal spreading are superimposed. The surface heat flow obtained from the stable geotherm model is 56.5 mW m-*, which seems to be a reasonable value, considering that the original Great Plains heat flow has been modified. Although the pre-rift history of crustal extension and magmatism has important consequences

events.

Attempts

thermal

record

by later

have been made to approximate

this transient temperature gradient. Reiter et al. (1986) assumed that the long period of pre-rift magmatism led to an equilibrium geotherm of 23.4’C from the surface to the base of a 35 km crust, and a mantle geotherm of 4.5’C, reaching 1200°C at 120 km depth. Morgan et al. (1986) apply a similar approach to pre-rift heating, assuming the base of the crust to be at the melting point, somewhat greater than 1100°C. Depending on the assumed crustal thickness, the gradients vary from about 25’C km-’ to 35’C km-’ and the mantle is assumed to have a gradient slightly

to this is

model shows that 15 into the lower 25 km of an

will give temperature

those we have assumed

in

anomaly. The pre-rift transients could have formed by lithospheric thinning as well as by magmatic injection. Within the rifting period, extension and magma injection is assumed to occur with a half-sinusoidal history, peaking ages on basaltic

at 25 Ma. A study of radiometric rocks in New Mexico (Aldrich et

al., 1986) shows an increase in the occurrence of intrusives up to 24-25 Ma, followed by an abrupt decline, consistent with the dike emplacement history assumed in our model (see e.g., Cook et al., 1979). The primitive magma intruded in the base

approximately

of the event has been erased

similar

prior to the inception of rifting, no inferences are made about the pre-rift history leading to this

much of the memory

and

to in-

less than 5% of the volume of the region. Although our model requires a thermal lithospheric anomaly

of the crust

the geologic

necessary

data, our results

our model. The pre-rift crustal thinning is 10 km, and the total amount of material injected is much

for the present-day thermal structure of the rift, this transient effect is difficult to quantify because from

3°C

25 km of the mantle

Plains structure

very similar.to

and

(Fig. 3). Although

is not

anomaly

A simple numerical

m.y. of injection

temperature

in the crust

(Irving

and

and Frey,

the upper

mantle

1984), assumed

150 km depth

on the intersection

and

of our initial

model

basanite

liquidus

1975;

DePaolo,

1979;

(Green, Ghiorso

is a basanite to originate 1450°C

geotherm 1973; and

at

based and a

Arculus,

Carmichael,

1980). The liquid is rapidly emplaced at the crust-mantle boundary at a superliquidus temperature (3 1450°C) by fracture propagation (Shaw, 1980) with minimal heat loss to the surrounding rock during

its ascent.

After emplacement,

as the

basanite rapidly loses heat to the surrounding country rock, liquidus phases pyroxene and spine1 begin to fractionate and latent heat is released. The crystallization is complete at the solidus temperature roxenite

(5 1200°C), and the newly formed dikes continue cooling conductively.

py-

382

G.Y.A.

The density crystallization (Kushiro,

to be 1500 J kg-’

1980)

a model

mineralogy

approximately

a latent

gg ’ (Bottinga

800 J

1978). Based on observations liths (Irving,

heat

for these

composites,

400

600 ‘,‘/

D.R

W1LLlAMS

(“C)

800 ,,I.

1000 1200 .,..,,I,_

1400

of the liquid and Bottinga, of 85% clino(Irving,

of fusion and P-T

of

Allegre,

of composite

1980) and the narrow

200

AND

’ based on heat

15% spine1 for the pyroxenites

we calculate

occurrence

"C-

for silicate liquids (Richet

1986). Using pyroxene,

Temperature

dike material upon be 3000 kg m-’

1980) and the heat capacity

is assumed capacities

of the injected is taken to

BUSSOD

xeno-

range of

the magma

is

modelled as being injected in the form of narrow dikes extending from a depth of 35 km to 55 km. The heat flow due to the injection is assumed to be distributed evenly within the region in one time step. The heat then conducts its surroundings.

out of the region

As the dikes are being

to

injected,

100

1

Fig. 4. Kilbourne

’

I

‘,

/

’ AfiJI

Hole thermal evolution

1

model-geotherm

re-

suits for 22 Ma, 20 Ma, 10 Ma, and the present, starting with the transient geotherm at 29 Ma (Fig. 3). The P-T-time of the original crust-mantle

path

boundary is represented by arrows.

the region is simultaneously extending by an amount equal to the volume of the dikes. The thermal history is not affected by whether dike injection

or lithospheric

extension

is assumed

to

be responsible for rifting. The crust directly above the region and for some distance to the sides is assumed

to extend

an equivalent

amount,

material to move upwards, carrying the geotherm with it, and providing the advection term in eqn. (1). The total extension in the intruded region is taken to be 50 km, so that the final width of the extended region is 200 km. This value for the present-day size of the extended and thinned area agrees well with the surficial fault distribution delineating the southern Rio Grande rift (see e.g., Chapin, 1979; Aldrich et al., 1986) and the seismic cross-section for this area (Sinno et al., 1986). The 200 km width represents 25% extension, in accord with the estimated value of 30% for the southern Rio Grande rift (Morgan et al., 1986). From our model, the mean extension velocity of the rift is 0.5 cm yr-‘, and the total resulting amount of thinning is approximately 9 km. the completion of extension and dike injection (20 Ma), the heat sources in the model comprise radiogenic heating and the heat flux from the mantle, and all heat is transported by conduction. At

it has reached

the present-day

structure.

and

therefore become progressively thinner. We have assumed the crustal thinning takes place over a region initially 150 km wide. This causes the

crustal

The geotherm is allowed to evolve under these conditions for 20 m.y., at which point we assume

Results Representative geotherms calculated by our finite difference model for various times using the parameters described above are shown in Fig. 4. Note that the maximum amplitude of the thermal “bulge” due to dike injection occurs at approximately 22 Ma. The bulge then relaxes conductively to the present-day feature of this model,

geotherm. An important distinguishing it from a

straightforward lithospheric thinning/asthenospheric upwelling model, is the development of an inverted temperature gradient in the upper mantle, as required by the piezothermometric data. The temperature variations with time below 60 km are very small, less than 100°C over the 25 m.y. duration shown. The maximum temperature reached at 60 km depth is 1069% 18 Ma. For comparison, the temperature at this depth 20 Ma was 1056*C, and the present-day temperature is 972°C. Conversely, the range in temperatures between 20 Ma and the present around 40 km depth is approximately 360% which corresponds to an

THERMAL

AND

KINEMATIC

average

cooling

complexes

present-day

estimated

OF THE

rate of 18”C/m.y.,

pine metamorphic The

MODEL

model

SOUTHERN

typical

(Dodson, surface

to be 82 mW m-*

and

is approximately

from

a maximum Note that km-‘.

is

the maximum at 19 Ma.

21’C

383

RIFT

Nathenson

and

noticeably

flow

average crustal temperature

ent for this model down

of al-

1976).

heat

model surface heat flow is 117 mW m-* The present-day

RIO GRANDE

gradikm-‘,

20 to 22 Ma of about the present-day upper-

m-*.

Guffanti,

This value

1988).

is

higher than our model result of 82 mW

Examination

of heat flow measurements

for

the SRGR (Reiter et al., 1986) reveals that the highest values (> 100 mW m-*) are concentrated along

the edges of the rift, while

surements vicinity

range

near-axis

mea-

from 70 to 90 mW m-*.

of Kilbourne

In the

Hole, a value of 80 mW m-*

is indicated

(Reiter

crustal thermal gradient is much steeper (dT/dZ = 25°C km-‘) and, as expected, near-surface

hydrothermal cause higher

circulation at the rift flanks would effective heat flow, but would not

gradients

have much effect on the deep thermal

35°C

are remnants

tures at depth.

Simply

of earlier extending

high

tempera-

this gradient

give misleading results for the geotherm and the time lag for thermal diffusivity taken into account when interpreting temperature gradients. The position

will

at depth, must be

near-surface and tempera-

ture of the original crust-mantle boundary through time is shown by the arrows (Fig. 4), the upward component

representing

crustal

thinning.

the rift (Lachenbruch. Morgan

be controlled

This model has significant implications for the present-day geologic structure as well as for the thermal structure of the rift. The final southern rift axial crustal (Fig. 2) consists

structure predicted by this model of a 22 km thick upper crust, a 4

km thick lower crust, tered by the magmatic composed of granulites,

which

is structurally

unal-

intrusives and principally and a 5 km thick lower

crust that is composed of 70% mafic intrusives 30% granulites. The mantle is divided into

and two

parts, an upper intruded region extending from 31 to 46 km composed of 70% spine1 pyroxenites (basaltic intrusive fractionates) and 30% compositionally re-enriched fertile lherzolitic mantle, and the region below 46 km, which consists of MORB lithospheric mantle material (Bussod and Irving, in prep.). The densities of these regions are also shown (Fig. 2). Studies of the present-day largescale structure of the rift provide an opportunity to test this model result against data.

actual geophysical

It is generally concluded that the average heat flow for the southern Rio Grande rift falls in the range of about 100 f 10 mW m-* (Sass et al., 1971; Blackwell, 1978; Reiter et al., 1979, 1986;

Concentrated

structure

of

and Sass, 1978). Seager and

also note

that

the variations

in

as well.

by the location

This

indicates

of the flank fractures

that

82 mW

m-*

may

represent a good estimate for the heat flow from the axis of the rift, the site of both Kilboume Hole our thermal

represent

structure

(1979)

1986).

heat flow within the rift may be due to localized near-surface magma bodies, an effect which may

and Present-day

et al.,

model.

a minimum

This

however

value

is taken

as the model

to

does

not take into account late-stage magmatism in the upper crust which would presumably cause an increase in surface heat flow at the present time. A large number of seismic studies have been done on the southern and Sanford (1976)

Rio Grande rift. Toppozada studied the refraction data

along a N-S line extending from the Gasbuggy shotpoint in northern New Mexico to the Texas border. They concluded that the crust at Albuquerque Applying buquerque

could

be

no

their

estimated

and

Kilboume

thinner 2”

than

dip

Hole

30

km.

between

Al-

constrains

the

crust of the southern Rio Grande rift to be thicker than about 20 km. The seismic evidence across the southern

Rio Grande

rift (McCullar

and

Smith-

son, 1977; Cook et al., 1979; Keller et al., 1979; Olsen et al., 1979; Sinno et al., 1986) is generally interpreted as consistent with asthenospheric welling beneath the Rio Grande rift, and

upthe

“anomalous mantle” P, velocities as evidence for a partially molten crust-mantle boundary. However, no evidence for high seismic attenuation is found at these depths (Davis, 1987), and an asthenosphere upwelling model cannot explain the piezothermometric data. We propose an alternate explanation, based on the similarities between

384

G Y.A. HUSSODAND

these southern Rio Grande cross-rift seismic results and the final state of our model. A discontinuity

at 20 to 27 km depth

(Cook

et al., 1979;

Keller et al., 1979; Sinno et al., 1986) is similar either

the lower-

model

to upper-crust

at 22 km,

and/or

assumed

to represent

Although

we have

matic

intrusion

physical crust, intruded

the 26 km boundary

constrained

crustal

altered

level, the actual

much more complex,

less distinction base.

zone.

the top of mag-

to be this single

a thermally

in our

the top of the intruded

case is undoubtedly

with somewhat

boundary

to

between lower

Using

the upper

crust,

seismic

and

an

reflection

studies of the central Rio Grande rift, Brown et al. (1979) concluded that the crust-mantle transition beneath the rift represents a complex zone of discontinuous reflector units. The boundary discovered seismically at 30-40 km depth is congruent with the interface between intruded lower

reviewed by Seager and Morgan (1979). The general consensus is that a high conductivity zone is present between about 20 and 30 km beneath the southern models

Rio

Grande

geomagnetic suggested

induction a N-S

the deep crust Colorado melting

rift,

have been proposed.

profile,

and upper

model.

intrusives

The

to 50 km

the axial line (7.95 km s-l)

and the cross-line

(7.7

km ss’) found by Sinno et al. (1986) may be due at least in part to anisotropy caused by the presence of dikes in the upper mantle oriented subparallel to the rift axis. Daggett et al. (1986)

studied

a large

set of

gravity measurements covering the southern Rio Grande rift. From these measurements, they constructed density distribution models for the structure beneath the rift. They interpreted these models as being consistent with anomalous mantle,

extending

these anomalies, and mantle

for from

with a simple

of N-S

assemblages

Hole lower crustal

trending

in the lower as Kilboume

xenoliths

are un-

and amorphous carbon and in intra- and inter-

granular fractures (Mathez et al., 1984; Dromgoole and Pasteris, 1987) which could result in anomalous electrical conductivities (Duba land, 1982; Frost et al.. 1989).

Results of the thermal on the piezothermometric

from 32 km (7.7 km s-‘)

(1970)

structure

crust could explain

usually rich in sulfides along grain boundaries

melting

based on a

Schmucker

mantle

presence

and high-grade

intruded area would not be sharply consistent with the observed gradual P-wave velocity

partial

to west Texas, inconsistent

Discussion

(8.3 km s-‘) (Sinno et al., 1986). Finally, the observed difference in P-wave velocity between

and

However,

two-dimensional

crust and intruded mantle at 31 km depth in our model. We assume the base of this intruded mantle layer to be at 46 km, but again the actual delineated, increase in

I, K WII.I.IAhlS

and

Shank-

model are superimposed arrays obtained from

both crustal and mantle xenoliths from Kilboume Hole presented earlier (Fig. 5). The apparent mismatch

between

the

crustal

and

mantle

arrays

strongly constrains the thermal evolutionary history beneath Kilboume Hole. The intrusive model fits the data well, given our set of geophysically reasonable parameter values. The crustal array temperatures representing “ peak” metamorphic matches the 20 Ma model temperatures. This result is consistent with the relatively slow cation diffusivities and reaction kinetics typical of these crustal rock mineral assemblages (Elphick et al.. 1985). Homogenization times for average size garnets from Kilboume Hole granulites are greater

elled by Daggett et al. (1986) is equal to our model column mass, indicating a similar gravity measure-

than 3 m.y. at 1000°C and 30 m.y. at 900°C. Given the range of model cooling rates (3 45 10°C/m.y.). closure temperatures are within the range of temperature estimates for the crustal garnet grant&es (78O”C-1000°C; Bussod and Irving, in prep.). Within error the mantle array

ment would be obtained. Keshet and Hermance (1986) have modelled the conductivity and magnetotelluric data for the SRGR. The electrical conductivity data is also

matches the present-day model geotherm well. Our calculations, however, do not take into account late stage ( < 20 Ma) crustal magmatism as proposed for this region (Morgan et al., 1986) which

related to magmatic activity at or near the base of the crust. This zone is described as transitional between normal upper mantle and lower crust. The total column mass to the 50 km limit mod-

THERMAL

AND

KINEMATIC

MODEL

OF THE

SOUTHERN

RIO GRANDE

TEMPERATURE (‘C)

RELATIVE PRECISION

80

Fig. 5. Model geotherm line), 10 Ma (stippled) on piezothermometric Model

basanite

results from Fig. 4, 20 Ma (dashed and present (solid line). superimposed

estimates (Fig. 1). Symbols as in Fig. 1.

liquidus

modified

from Green

(1973)

and

Arculus (1975).

may explain why our model present-day geotherm falls below the uppermost mantle thermobarometric estimates (Fig. 5). An initial transient crustal geotherm of 25°C km-’ (30 Ma) also improves the fit; however, the data does not warrant these modifications. In addition, the olivine-spine1 thermometric estimates for the granuloblastic rocks (Fig. 5), which represent a relic thermal perturbation (2 10 Ma), are in agreement with a 20-22 Ma model geothermal perturbation at the crustmantle boundary. Maximum thermal perturbations within this zone occurred 22 Ma at a depth of 41 km, within the modified upper mantle. This thermal maximum is within 30°C of the “dry” melting temperature of the MORB-related Kilboume Hole mantle spine1 lherzolites (Takahashi, 1986; Bussod, 1991), and is above the solidus temperature of the lowermost crustal assemblage (Irving, 1974; Vielzeuf and Holloway, 1988). This implies pervasive recrystallization (T/T, > 0.95), as well as possible melting and chemical re-equilibration at the crust-mantle boundary at that time (Bussod and Irving, in prep.). The narrow P-T range for composite xenoliths and the

R1l-V

385

deepest protogranular mantle samples which show no thermal perturbation are also consistent with our thermal model and thus limit the zone of 70% magma injection to a present-day depth of approximately 26 to 46 km. The granuloblastic textures of the crustal and uppermost mantle xenoliths are interpreted as textural evidence for our model which predicts pervasive recrystallization at and near the crustmantle boundary during the first 10 Ma of rifting due to a thermal perturbation. Isotopic data on one of the shallowest granuloblastic mantle xenoliths (KH77-7) indicates this sample is in isotopic equilibrium with the pyroxenites found in composite mantle xenoliths (Roden et al., 1988). Furthermore, this sample has been shown to be chemically re-enriched in a basaltic component (Irving, 1980; BVSP, 1981; Irving and Frey, 1984; Roden et al., 1988). These observations, along with the modelled maximum temperatures, are taken to imply that the uppermost mantle and lowermost crust were re-enriched in a basaltic component between 20 Ma and 30 Ma due to magmatic intrusion. Mineral isotopic data of a lower crustal granulite from Kilboume Hole yields a re-equilibration age of 34 f 10 Ma, whereas bulk isotopic model ages of the lower crust indicate 1.6 f 0.1 Ga (Padovani and Reid, 1989) again consistent with our model. The deepest granuloblastic mantle xenoliths (KH77-15, KH77-11, KH77-12; open circles, Fig. 5) fall in the region where near-solidus temperatures are never attained in our model. These xenoliths do not show the extreme compositional re-enrichment of the shallowest granuloblastic and composite mantle samples, and are in isotopic disequilibrium with the pyroxenite intrusives (Roden et al., 1988). Unlike the granuloblastic samples, the protogranular mantle xenoliths are compositionally more barren and do not show the progressive bulk Fe enrichment or increasing LREE with decreasing depth (Bussod, 1982; Bussod and Irving, in prep.). This implies that the degree of basaltic re-enrichment of the uppermost mantle may be temporally linked to the intrusive event 20-30 Ma as in the case of the textural re-equilibration. Evidence for partial melting of the wall rock can be observed in some composite xenoliths

386

BUSSOI) AND

ti.Y.A.

(Bussod upper

and Irving, mantle

although

this

pyroxenite the

is contingent mineral

fabric

pyroxenite

shape

along

within

the

age of the

flow continued amphibole,

anisotropy

beyond

15) and rare mica-bearing

possibly

and

deforma-

dikes indicate

the solidification

The presence

amphibole

in

wall rock to the

with plastic

of dissemi-

in veinlets

samples

granuloblastic and composite that a hydrous liquid phase recrystallization, complex.

Continuity

the pyroxenite

time for the intrusives. nated

to 30 Ma,

on

in the lherzolitic

interface

tion features plastic

that a fertile

prior

dikes, which is unknown.

observed

tectonic

1990) suggesting

may have existed

(KH77-

present

only in

xenoliths, suggests was present during

related

to the intrusive

lower crust and upper thinning (25% extension) SRGR

characteristics.

a process, position

and

reasonably profile

It seems plausible could

type

of this

constrained.

However, spheric

lithospheric

cause of the thermal with the southern magmatic

thinning anomaly

Rio Grande

intrusion

event

are

temperature and/or

for the pre-rift

are precluded

to a

magnitude,

intrusive

thinning

is required

upwelling

The

A transient

due to lithospheric

matic intrusion

that such

lead naturally

structure.

timing

WILLIAMS

mantle produces crustal and matches present-day

if continued,

spreading-center

L).R

and

maghistory.

astheno-

as the primary

currently

associated

rift. The timing

at the crust-mantle

for

boundary

beneath Kilbourne Hole is consistent with a regional event associated with the origination of a

These observations are all consistent with our model and further support the concept that the basaltic re-enrichment is correlated with a crust-

continental extension zone and the formation of a granulite metamorphic complex in the lower crust. The modelled present-day heat flow of 82 mW

mantle

mm2 is in agreement with present-day heat flow measurements in the vicinity of Kilbourne Hole. The recent emplacement age of the xenolith suite

boundary

intrusive

complex.

Other

than

the mantle composites, candidates for such a complex in the crust and mantle are the ortho- and clino-pyroxenites

described

by Padovani

(1977)

Baldridge (1979) Irving (1980), Dromgoole and others. Although barometric estimates

(1984) cannot

at Kilbourne Hole maar ( < 100 ka) precludes other modern regional thermal or structural turbation

of the lower

crust

and

upper

any per-

mantle.

their place

This model predicts that the anomalous compressional wave velocities measured in this region are

them in the range of 25 to 50 km depth 20 Ma ago, using our thermal model or within the present depth range of the intrusive zone in our model

not the result of a partially molten crust-mantle boundary but are due to a relatively young (< 20 Ma) N-S trending mafic intrusive complex. This is consistent with observed seismic anisotropies in

be made accurately closure temperatures

(26-46

for these assemblages, (900-1050°C) would

km).

the vicinity of the crust-mantle boundary in the SRGR (Sinno et al., 1986) and the unusually large

Conclusions Evaluation

of the P-T-time

mantle boundary complex thermal last

30 m.y..

at Kilbourne and kinematic

The

path of the crustHole reveals a history over the

piezothermometric

data

Kilbourne Hole xenoliths support the idea inverted temperature gradient in the upper existed within the past 20 m.y.. This is sistent with a simple lithospheric thinning

for

that an mantle inconmodel

by asthenospheric upwelling, but rather requires a rapid transfer of heat directly to the crust-mantle boundary from a deeper mantle source. Numerical thermal calculations constrained by piezothermometric data at depth show that a model of regional crustal underplating by intrusion of magma in the

proportion of composite and mafic lithologies found in the xenolith popuiations of this area (Padovani and Carter, 1977; Fodor, 1978; Baldridge,

1979; Irving,

others). The extent ble to the northern

1980; Dromgoole,

1984; and

to which this model is applicaand central Rio Grande rift is

uncertain at this time as the original rift structure in these regions may have been modified by late stage volcanism associated with large central volcanic centers unrelated to rift development. Acknowledgements The authors would like to thank J.-C. Mareschal, L. Hirsch and an anonymous reviewer

THERMAL

AND

KINEMATIC

for their careful

MODEL

reviews

G.B. acknowledges

OF THE

SOUTHERN

and critical

RIO GRANDE

Brown,

comments.

of Geophysics

R.L.,

Krumhansi,

P.A., Chapin,

Cook, F.A., Kaufman,

the support of T. Shankland

and the Institute

387

RIFT

and Planetary

Physics at Los Alamos.

CGCORP

seismic reflection

In: R.E. Riecker Magmatism. Bussod,

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