Crustal structure of the Northwestern Basin and Range Province from

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JOURNAL

OF GEOPHYSICAL

RESEARCH,

VOL. 95, NO. B13, PAGES 21,823-21,842, DECEMBER

10, 1990

Crustal Structure of the Northwestern Basin and Range Province From the 1986 Program for Array Seismic Studies of the Continental Lithosphere Seismic Experiment HARLEY

M.

BENZ

U.S. Geological Survey, Menlo Park, California ROBERT B. SMITH

Department of Geology and Geophysics,University of Utah, Salt Lake City WALTER

D. MOONEY

U.S. Geological Survey, Menlo Park, California The Basin and Range province of the western United States is characterized by active east-west

extension at strainratesof--•1 cmyr-1 , highregional heatflowof--•90mWm-2, andwidespread late Cenozoic bimodal rhyolitic-basaltic magmatism. The 1986 Program for Array Seismic Studies of the Continental Lithosphere Basin and Range seismic experiment imaged a portion of northwestern Nevada to determine the crustal structureand to assessreported differencesbetween refraction versus reflection determinations of Moho depth and how the crustal composition and structure has been influenced by volcanic and extension mechanisms.Our interpretation of the refraction/wide-angle reflection data suggeststhat the crust is fairly uniform in thicknessand varies by less than 5 km over the 280 km east-westprofile and 3 km over its 220 km north-southlength. We characterize the velocity structure by five layers: (1) an upper most crust, composedof sedimentary rocks and basement that

hasan averagevelocityof 5.7 km s-i , (2) a middlecrustthatextendsto a depthof 18-22withan average velocityof6.1kms-1 , (3)a 10-12kmthicklowercrustwithanaverage velocityof6.6kms-1 , (4)a 2-5kmthicktransitional crust-mantle boundary defined bya 7.6kms-l velocity,and(5)anupper mantlewithanaverage Pnvelocityof 7.9-8.0kms-1 . Thesedatashowa uniformily thickcrustof 28 km +5 km beneath northwestern Nevada and a lack of evidence for significantlow-velocity crustal layers. In general, the entire crustal column exhibits uniform positive velocity gradients of 0.02-0.04

s- . Themidcrustal discontinuity is interpreted asa smallvelocityincrease of 0.2km s- l at --•18km depth. The absenceof a significantmidcrustalreflectionand the presenceof large velocity gradients throughout the crust arises either from (1) a systematicincrease in mafic composition with depth, or (2) possibly high lower crustal pore pressure that acts to decrease the velocity contrast across the midcrustal boundary. Our interpreted seismic velocity structure is consistent with a broad range of bulk compositionsfor the crust. Upper crustal velocities are consistentwith bulk compositionsthat range from metasedimentsand granites at shallow depth to diorites at midcrustal levels, while lower crustal velocities suggestbulk compositionsthat range from diorites and amphibolities, at midcrustal depths, to mafic granulites and gabbrosat the base of the crust. The uniform crustal thickness of 30-34 km in northwestern Nevada is closeto the crustal averagefor the Basin and Range province but is thin when comparedto surroundingregions, suchas the Sierra Nevada and Colorado Plateau where crustal

thickness isgreater than40kin.Similarily, northwestern NevadaPnvelocities of 7.9-8.0kms-1 are neartheprovince-wide average of 7.9kms-l . Temperature-corrected Pn velocities, froma 90 mW m-2 geotherm fortheBasinandRangeto 60mWm-2 forthestableinteriorof theUnitedStates,are withinonestandard deviation of thecontinental average of 8.05kms-i andsuggest a uniformupper mantle compositionacross the Basin and Range. In addition to these observations, the homogeneity of the velocity structure beneath the western Basin and Range argues for a youthful Moho and crust that has been reworked by province-wide late Cenozoic extension, episodic magmatism, and underplating.

INTRODUCTION

Since the King Survey of the 40øN parallel (1867-1872), the Basin and Range province of the western United States has been the site of numerous geologic and geophysical studiesaimedat understandingthe geologicprocessesof this tectonically active region. A broad zone of continental extension since Miocene time, the northwest Basin and

Range province consists of a series of northeast trending

subparallelmountain ranges, spacedat roughly 25-km interval, separatedby alluvial-filled basins. Covering an 800,000

km2 area,theBasinandRangeischaracterized byrelatively high topography (averaging 1250 m above sea level), high

regionalheat flow of--•90 mW m-2, a historyof late Cenozoic silicic-basaltic volcanism, Quaternary normal faulting, and historic earthquakeactivity exceedingmagnitude 7.

Paper number 90JB01537.

Previous refraction surveys (Figure 1) of the western Basin and Range [Pakiser, 1963; Eaton, 1963; Hill and Pakiser, 1966; Prodehl, 1979] demonstrated that the crust is

0148-0227/90/90JB-01537505.00

thin

Copyright 1990 by the American Geophysical Union.

21,823

in contrast

to the

thicker

crust

beneath

the

Sierra

21,824

BENZ ET AL.' BASINAND RANGECRUSTALSTRUCTURE

120ø

115ø

110ø

105ø

IDAHO

OREGON WYOMING River I

!

ß ....

B

;or

F

SP11

"1

oUTA• b



40ø

COLORADO

Colorado

NEVADA

Plateau [ NEW MEXICO

ARIZONA

35ø

0 100 200300 I

I

I

KM

I

Fig. 1. Index map of Basin and Range showinglocation of the 1986PASSCAL Basin and Range seismicexperiment (profiles C and D). Other profiles discussed are A, Shasta-Mono Lake, Sierra Nevada [Prodehl, 1979]; B, Eureka (Eur)-Fallon (Fal), western Nevada [Thompson et al., 1989]; C, northwest Nevada, this study; D, northwest Nevada, this study; E, Boise-Eureka, central Nevada [Hill and Pakiser, 1966]; F, Delta-Eureka, western Utah [Keller et al., 1975]; G, central Utah (a, Keller et al. [1975]; b, Braile et al. [1974]; and H, Colorado Plateau [Roller, 1965].

Nevada

on the west and Colorado

Plateau

on the east. The

seismicexperiment and permits a joint interpretation of the two types of data. Our interpretation of the 1986PASSCAL Basin and Range seismic experiment focuses on developing a model of crust and upper mantle structurebasedon iterative amplitude and travel time modeling of selected seismic refraction/wideangle reflection data that we felt was representativeof the deep crustal structure of northwesternNevada. From initial 7.6 and7.9 km s-l . Deepseismic reflection profiling across comparisonsof individual record sections and synthetic northern Nevada [Allmendinger et al., 1987; Hague et al., seismogram modeling it is apparent that generalized one1987] reveals (1) a moderately reflective upper crust with dimensional velocity models, determined for each shot evidence of planar to listric normal faults and asymmetric point, are adequate to assess the regional nature of the basins, (2) a highly reflective lower crust marked by subhor- crustal structure of this region. The justification is detailed izontal, discontinuousreflections, and (3) a laterally discon- later in this paper. Our discussion complements several tinuous but moderately reflective Moho with approximately ongoing studies of the 1986 PASSCAL seismic data includhalf the vertical relief determined from earlier refraction ing (1) two-dimensional travel time and ray theoretical synthetic seismogram modeling of Catchings and Mooney models of the Moho [Klemperer et al., 1986]. While previous reflection and refraction work addressed [1990] and Holbrook [1990], (2) tau-p inversion [Matulevich, fundamental questions about Basin and Range structure, 1989; Hawman et al., 1990], and (3) two-dimensional inverprimary characteristicsof the crustal structure of northwestern Nevada, based on previous surveys, are (1) eastward crustal thickening from ---25to ---30km toward the interior of the Basin and Range and thickening westward to ---50 km beneath the Sierra Nevada, (2) a simple velocity structure without widespread evidence of crustal low-velocity zones, and (3) anomalouslylow upper mantle Pn velocitiesbetween

individual

studies did not have sufficient

resolution

to ad-

dressdetailed issuesof Basin and Range crustal composition and tectonics. The 1986 Program for Array Seismic Studies of the Continental Lithosphere (PASSCAL) seismicexperiment was designed to insure sufficient resolution of refraction and wide-angle reflections from throughout the crust by using shorter seismometer spacings and a multiplicity of shots. In addition, the Consortium for Continental Reflection Profiling (COCORP) 40øN reflection profile [Hague et al., 1987] is coincident with portions of the 1986 PASSCAL

sion of travel times [Braun, 1989]. WESTERN

BASIN AND RANGE GEOPHYSICS OVERVIEW

Geophysicalsurveysand geologicinterpretationsof Basin and Range structure and tectonics fill a large body of literature. Reviews of the regionalgeologyand geophysicsof the Basin and Range include those of Thompson and Burke [1974], Smith [1978], Eaton [1979], Speed [1982], Allmendinger et al. [1987], Pakiser [1989], Smith et al. [1989], and

BENZ ET AL.' BASIN AND RANGE CRUSTAL STRUCTURE

Thompson et al. [1989]. The reader is referred to R. D. Catchings and W. D. Mooney (Basin and Range Crustal and Upper Mantle Structure Along the 40øN Parallel, Northwest Nevada, submitted to Journal of Geophysical Research, 1990) for a description of the regional geology and tectonics. We will focus our attention on geophysical surveys in the western Basin and Range that are pertinent to this study. In 1986 PASSCAL Basin and Range seismic experiment [PASSCAL Working Group, 1988] transects an area affected by Mesozoic and early Cenozoic eastward overthrusting, followed by extension beginning in the middle Cenozoic. The high heat flow, extensive Quaternary normal faulting, and volcanic rocks observed along the transect, are primarily the result of the late Tertiary extensional tectonics. What is unclear is how the crustal structure of previous tectonic and volcanic episodes, principally Mesozoic and Cenozoic compression and late Cenozoic silicic-basaltic volcanism, has been changed by younger extension. In the mid-1960s, interpretation of regional seismicrefraction profiles recorded in the Basin and Range by the U.S. Geological Survey (Figure 1) indicated a thin (22-32 km) crust overlying an upper mantle with P wave velocities

21,825

[Borcherdt et al., 1985] using vertical, 2-Hz geophones. The experiment included an industry standard 396-channel, 25km-long reflection spread, with ---60-m station spacings, located near the center of the experiment at the crossing of the two long profiles. Both profiles were reversed using charge weights of 400-2700 kg for shots approximately 40 km apart. A total of 28 shots were detonated. Whitman and Catchings [1987] provide details on the station deployment, shot size, instrumentation, and other pertinent information related to this experiment. North-south deployment. Along the north-south profile recorders were spaced at ---1.5-km intervals from shot point (SP) 8 near Winnemucca, Nevada, to the Carson Sink, shot point 11 (Figure 2). The coincident reflection spread was deployed near the center of the north-south profile and north of SP4 in Buena Vista Valley. The profile trends subparallel to the basins and ranges (Figure 2) and covers a total distance of 220 km. Five shots, with an average shot spacing of 40 km, were recorded across the north-south array, as was a fan shot from

SP1.

Throughout this paper individual profiles and record sections will be described by SP (shot point) followed by the between7.6 and7.8 km s-1. Prodehl's[1979]reinterpreta- corresponding shot point number. For example, the refraction and summary of these data suggesteda 29-35 km crust tion record section for shot point 11 on the north-south witha 7.9km s-1 uppermantlevelocity.StauberandBoore profile will be referred to as SP11. Only a brief description of [1978] used quarry blasts and nuclear explosions from the the field parameters germane to further discussionsin this Nevada Test Site (NTS) in southern Nevada to interpret paper are presented below. unreversed refraction profiles in northern Nevada. They Interpretations of the refraction data from the end shot suggestan anomalously thin crust and a low upper mantle points, SP8 and SP11, are presented in Figure 3. These two velocityof7.8kms-1. Priestley et al. [1982]concluded from record sectionsare generally representative of the refraction an independent north-south refraction experiment, using data recorded along the north-south profile, and they are Nevada Test Site (NTS) nuclear explosions and quarry sufficient for resolving the crustal structure beneath Grass shots, that the crust is as thin as 20 km in northwestern Valley and Buena Vista Valley. Nevada. All of these studies emphasized an anomalously East-west deployment. The east-west profile, trending low upper mantle velocity and thin crust. perpendicular to the structural grain, was constructed from Basin and Range heat flow measurements reported by two 140-km deployments (Figure 2). The western deployBlackwell [1978] and Lachenbruch and Sass [1978] show ment extended from SP1, near Gerlach, Nevada, to SP4, and regionalheatflowvariations of 85-120mW m-2 fromsouth the eastern deployment continued from SP4 to SP7, near to north across the 1986 PASSCAL survey area. Eddington Eureka, Nevada. Multiple shots from shot points 1, 3, 4, 5, et al. [1987] determined regional extension rates by summing and 7 provided the coverage necessary to produce the 280 seismic moments of historic earthquakes and deduced an km compositerefraction/wide-angle profile. Shots from shot east-west extension rateof---1cmy-1 forthenorthern Basin points 2 and 6 provided additional data on upper and and Range. This rate is consistent with models by Lachenmidcrustal structure. Coincident reflection spreads were bruchandSass[1978]of 5-10mmy-1 extension onthebasis deployed in Buena Vista valley during the western deployof observed heat flow and suggests several responsible ment and in Dixie valley during the eastern deployments tectonic mechanisms including magmatic intrusion, mag- (Figure 2). From the composite refraction profile, we have matic underplating of the crust and lithospheric stretching. emphasized modeling and interpretation of the data from shot points 1 and 7 in this study. DESIGN

OF THE 1986 PASSCAL

AND RANGE

SEISMIC

BASIN INTERPRETATION

EXPERIMENT

The 1986 PASSCAL Basin and Range seismic experiment was principally designed(1) to image the crustal structure of an enigmatic region of the western Basin and Range, (2) to measure accurately the underlying upper-mantle velocity structure, and (3) to address apparent differences between wide-angle refraction and near-vertical reflection data. To achieve these objectives a 280-km-long east-west refraction/ wide-angle reflection profile and a 220-km long north-south profile were recorded during the 1986 PASSCAL seismic experiment (Figure 2). Both profiles were recorded on 120 vertical-component cassette recorders [Healy et al., 1982] and 26 three-component GEOS digital instruments

AND VELOCITIES

OF CRUSTAL P WAVE

PHASES

Based upon the interpretations of Eaton [1963], Stauber and Boore [1978], and Priestley et al. [1982], western Nevada crustal

structure

showed

lateral

variations

in crustal

thicknesses,dominated by crustal thinning in the vicinity of our study area. However, during preliminary modelingof the 1986PASSCAL Basin and Range seismicdata, in particular, the profiles SP8-SP11 (Figure 3) and SP1-SP4 (Figures 4 and

5) we notedthatthePa, PmP, andPn branches hadsimilar apparent velocities. In fact, travel times and amplitudes of individual phases from reversed shots matched within the expected errors due to near-surface site geology. These

21,826

BENZET AL.' BASINANDRANGECRUSTAL STRUCTURE

41 ø N

SP8* ALLUVIUM

Winnemuccaß :.

VOL CANIC ROCKS GRANODIORITE

GABBRO

Gerlach

GRANITE

ß

CLASTIC SEDIMENTARY ROCKS LIMESTONE

SPl'k.

QUARTZITE

- NORMAL FAULT

THRUST FAULT SHOTPOINT LOCATION REFRACTION STATION LOCATION REFLECTION ARRAY LOCATION

Lovelock

40 ø N

SP10'

Fallonß o

Austinß

20

km i

•o w



w

•o w

Fig.2. Location mapof the1986PASSCAL Basin andRange seismic transect. Circles represent theprincipal PASSCAL shotpoints (SP)referred tointhediscussion. Theheavy linesrefers tothecoincident reflection arrays. The COCORP 40øN profile (notshown) wascoincident witheast-west linefromSP4toSP6andparallel totheprofile, --•15

km south,fromSP1to SP3.IndexmapfromMatulevich[1989].

similarities suggestedthat data modeling via a one- source. In the time term correction, a static travel time shift dimensionalsyntheticseismogram methodwasviable.How-

is appliedto each seismogramin a record sectionsuchthat it

ever, we also noted differences in the travel times and

is correctedto the elevationof the respectiveshot. Shot

amplitudes of the phaseson the SP7profile,suggesting that point elevationsrange between about 1200and 1900 m. On SP7wasanomalous compared to theotherprofilesacquired the north-south profile,the time term corrections averaged in the experiment.To facilitatea description of the seismic -0.08 s and rangedfrom -0.46 s to +0.25 s, whereasthe data, the discussionwill be subdivided between the north- timetermcorrections for the east-westprofileaveraged0.02 southprofile (SP8 and SPll; Figure 3) and the east-west s and ranged from -0.26 s to +0.43 s. profile(SP1, SP7, and SP4;Figures4 and 5). North-south profile. Alongthenorth-south profile,upper In this paper,emphasisis placedon amplitudeand travel and lower crustal P wave phaseswere well recorded and time modeling of the seismicrefraction data from both the

correlatable overtensof kilometers. ThePaphase isprom-

north-southand east-westprofiles to estimatedepth- inently observed on SP8 and SPll as first arrivals out to dependentvelocity structure,structureof the crust-mantle distances of 110-120km (Figure3). From 5 to 30 km, the P•

boundary,and seismicattenuation (Q-•). Recordsections apparent velocity increases rapidly from5.0to5.5kms-• are plottedtracenormalizedwith a reducingvelocityof 6.0 and then to5.8kms-• at60km.Beyond 60kmthePaphase kms-•. In Figures 3, 4, and5, eachtracewasband-pass increases only slightlyto about6.10 km s-• (Figure3). A

filteredbetween0.5 and 12.0Hz, normalizedrelative to the large amplitude phase fordistances lessthan60km,thePa maximumtraceamplitudewithinthe plottingwindow,and amplitudedecreasessignificantlybeyond60 km. Overall,the plottedto a specifiedwidth.Additionally,all recordsections timeterm corrected Pa phasedoesnot exhibitthe large were corrected for systematic travel time variations due to traveltimevariationsindicativeof significant lateralvelocity topography andlateralvariations in thickness andvelocityin variations,but it does suggesta moderateincreasein velocthe sedimentarycover usingthe time term algorithmde- ity with depth. scribedby Kohler[1988].Time-termcorrections wereonly A relatively weak midcrustalwide-anglereflection, de-

applied to stations thatclearlyrecorded Pa phases from notedasPiP,wasobserved following thePaphaseonthe

three or more shots. The time term correction thus reduces the data to a datum coincident with the elevation of the

data from SP8 between50 and 110 km (Figure 3). The amplitudeof this midcrustalreflectionvaries significantly

BENZ ET AL.' BASIN AND RANGE CRUSTALSTRUCTURE S SP8

21,827 N

TRACE NORMALIZED

6-

-

6

5

-

5

-

ti

-

3

-

2

•-

1

L o

r- 0

----1

-1 I

-2_ 200

I

-180



-160

I

I

-lti0

-120

I

-100

I

-80

I

-60

I

-ti0

-2 0

-20

SP11

6

-

6

-

5

-

-2

I 0

20

I

I

I

I

I

ti0

60

80

100

120

DISTRNCE

-

3

-

2

-

!

-

0

-2

I I ti0

160

180

200

(KH!

Fig. 3. Tracenormalized seismic datarecorded fromSP8andSP11onthenorth-south profile.Eachtracehasbeen band-pass filteredbetween 0.5and12.0Hz, timetermcorrected to theelevation of thecorresponding shotpoint,and

plotted using a reducing velocity of6.0kms-• . Stations northoftheshotpointarerepresented bypositive distances. Keyto phaseidentifications (usedhereandin thefollowing figures): Pg,thedivingor continuously refracted P wave in the crystalline uppercrust;PiP, the wide-angle reflection fromthe midcrustal discontinuity; PmP,thewide-angle reflectionoffthecrust-mantle boundary(Moho);andPn, theuppermantleheadwave.Predicted traveltimecurvesfrom therespective velocitydepthfunctionareshownasdarklines.Precritical reflections areindicated by dashed lines.

w

sP1

TRACE

NORMALIZED

E

5 q

'

1

i

r- o -1

-1 I

o

20

q0

60

80

100

120

lq0

I

I

160

180

I

200

220

2qO

TRACE

$P7

260

280

NORMALIZED



1

-2

-280

I

I

I

I

I

I

I

I

I

I

I

I

i

-260

-2qO

-220

-200

-180

-160

-lqO

-120

-100

-80

-60

-qO

-20

DISTRNCE

i -2 0

(KH)

Fig. 4. Tracenormalized seismic datarecorded fromSP1andSP7ontheeast-west profile.InsetonplotSP1data showsPn phaseat an expandedscale.SeeFigure3 for detailsof plot.

21,828

BENZ ET AL ' BASIN AND RANGE CRUSTAL STRUCTURE W

SP4

TRACE

•< 3 ,

:

NORMALIZED

E

"'

3

2

•-

2

I

!

0

0

-I

I

I

-1110

!

-120

-100

I

I

-80

-60

I -rio

I

I

I

I

I

I

I

I

I

-20

0

20

•t0

60

80

100

120

lttO

DISTRNCE

Fig. 5.

-!

(KH}

Trace normalized seismicdata recorded from SP4 on the east-west profile. The similarity of travel times and

amplitudes of theprinciplephases(Pa,PiP, andPmP)indicates a generally uniformcrustalstructure.SeeFigure3 for details of plot.

between shot points on both the north-south and east-west profiles. We interpret the observed PiP phase as a largeamplitude phase only on the record section from SP9 (not included here) on the north-south profile (R. D. Catchings and W. D. Mooney, submittedmanuscript, 1990). The travel time and amplitude character of the midcrustal reflection changeslaterally over tens of kilometers suggestinglocalized vertical and horizontal changesin midcrustal velocity structure or changesin local surfacevelocity structure. The lower crustalrefraction or diving phase,Pi, was not well recorded or not easily recognized. Observed clearly on both SP8 and SP11 profiles, PmP is a relatively large-amplitude secondary arrival, first evident between 70 and 80 km at about 3.0 s (Figure 3). The critical distance for this phase is between 85 and 95 km. The mantle head wave, Pn, has an unusually weak amplitude that is generally observed as a first arrival at distancesgreater than 140 km. On SP8 and SPll, the estimated apparent velocity of

thisphaseis approximately 7.9 km s-I .

the record section and is seen from 35 to 180 km. As detailed

in the next section, we believe this unique feature is due to a thin, high-velocity layer in the mid crust. The lack of a prominent midcrustal reflection on most of the north-south and east-west profiles strongly argues against an abrupt compositionalboundary such as might be expected between granitic upper crustal and gabbroic lower crustal rocks. The PmP reflection is the predominant seismic phase on all east-west record sectionsfor distancesgreater than about 70 km, except on SP7. From SP7, PmP is first observed at about 80 km and 3.5 s and is a weak phase compared to PiP. Beyond 220 km (SP1, Figure 4) the PmP travel time branch

asymptotically approaches -6.7 km s-•. ThePn phaseis seen on SP1 as a weak phase propagating with an apparent

velocityof about8.0 km s-1 (seeinset,Figure4). The refraction data from SP4 (Figure 5) shows, in general, travel time and amplitude variation with distance similar to that observed from SP1 (Figure 4). The similarity of travel

timeandamplitudes of the phasesP•, PiP, andPmP,noted

in either direction from SP4, suggeststhat the crust near SP4 is laterally homogeneous.Thus we suggestthat the data from both the north-south and east-west profiles indicate a laterally homogeneouscrustal structure in northwestern Nevada, Indeed,whiletheP• phasedisplays generally uniformtravel with the largest anomaly being an unusually strong PiP time and amplitude behavior with distancein the north-south reflection from SP7 at the east end of the study area. record sections, SP8 and SP11, it exhibits a larger degree of variability in both travel time and amplitudeon the east-west AMPLITUDE AND TRAVEL TIME MODELING sections, SP1, SP7, and SP4 (Figures 4 and 5). We attribute this variability to a combination of receiver site (nearSyntheticseismogram modelingof the crustalphasesPa, surface) effects and complexities in the uppermost crustal PiP, PmP, and the mantle head wave, Pn, were used to velocitystructure.The largestvariationsin P• traveltime determine the depth-dependent velocity and attentuation East-west profile. The east-west profile crossesthe geologic strike of the basins and ranges of this region and, as a result, it might be expected that the correlation of seismic phases would be more difficult to identify on this profile.

andamplitudeare bestseenin SP7(Figure4) wherethe P• amplitude decreases significantly beyond 40 km. We interpret the noticeable travel time delay of approximately 0.5 s beyond 40 km as a shallow upper crustal low-velocity zone, the details of which will be presented in the following

(Q-l) structureof northwestern Nevada. In this paper, emphasisis placed on modeling observed amplitude varia-

tionsof thePa andPmP phasesthat arewell recorded from

all shots.The PiP and Pn phasesare intermittently observed and weak compared to PmP, yet they provide important section. information on the crust and mantle velocity structure. Not TheP• phaseobservedon SP1is similarto thatobserved observed as a first arrival and difficult to consistently idenon the north-south profile and is characterized by increasing tify, the lower crustalturning phase, Pi, is difficultto model apparent velocity with offset, implying an increase in veloc- and will not be addressed in this study. Because we reity with depth.In general,P• velocitiesrangingfrom5.0 km stricted our synthetic seismogram modeling to use of the s-• in theshallowcrystalline uppercrustto 6.1kms-• near one-dimensional reflectivity method, particular attention the midcrustal boundary. The midcrustal reflection, PiP, is a was placed on modeling the depth-dependent velocity strucrelatively low-amplitude phase typically observed between ture and seismicattenuation. It is difficult to separateintrin60 and 120 km at about 1.5 s. In contrast, the middle crustal sic seismic attenuation from scattering attenuation [Aki, phase (PIP), noted from SP7, is the most prominent phase in 1980; Frankel and Wennerberg, 1987], and we did not

BENZ ET AL.: BASIN AND RANGE CRUSTAL STRUCTURE TABLE

z, km

a, kms-l

la.

North-South

/3,kms-1

0.00 1.50 1.50 5.00 9.50 18.50 18.50 29.50 31.00 33.60 34.50 40.00

2.80 3.50 5.40 5.95 6.05 6.22 6.40 6.85 7.65 7.70 7.90 7.92

SP8* 1.62 2.02 3.12 3.43 3.49 3.59 3.70 3.93 4.42 4.45 4.56 4.59

0.00 1.60 1.60 6.00 9.50 17.00 28.00 29.00 33.50 34.50 40.00

3.00 3.50 5.50 6.00 6.10 6.25 6.85 7.50 7.60 7.85 7.90

1.73 2.02 3.18 3.46 3.52 3.61 3.95 4.33 4.39 4.53 4.56

Profile

P,gcm-3

Q-

Qt•

1.31 1.58 2.30 2.51 2.54 2.61 2.68 2.87 3.15 3.17 3.25 3.27

200 200 200 200 400 400 1000 1000 1000 1000 1000 1000

100 100 100 100 200 200 500 500 500 500 500 500

1.39 1.58 2.34 2.53 2.56 2.62 2.85 3.09 3.13 3.23 3.25

200 200 200 200 400 400 1000 1000 1000 1000 1000

100 100 100 100 200 200 500 500 500 500 500

SPlit

21,829

reference frequency was chosen as 5 Hz, and the reference velocities are listed in Table 1. A frequency-dependent, complex velocity law was chosen to avoid introducing acausal body wave arrivals common with frequencyindependent complex velocity laws [Mt;iller, 1985]. The reliability of the interpreted velocity-depth functions will be illustrated by two comparisons. First, a comparison of the observed and synthetic seismogramsfor each shot will emphasize the match between travel times, relative amplitudes, and frequency variations of individual crustal P wave phases. For seismic data acquired over a large geographic area, such as ours, especially where site-dependent amplitude variations can be large and amplitude modeling is difficult. Therefore we favor a combination of true amplitude and normalized plots to match the general amplitude features, such as position of critical reflections and relative amplitudes of correlatable P wave phases within a trace. When modeling relative amplitudes useful comparisons are

Pg/PmP and Pn/PmP ratios, even when they are only observable over a few tens of kilometers. For example, the

P• and PmP phases,for mostshots,are observedon the same trace from approximately 80 to 120 km. We feel that amplitude modeling using true amplitude and normalized record sections provide a reliable determination of the velocity structure in northwestern Nevada. Because we have choosen to forward model the observed data, no formal

z, depth' a, P wave velocity;/5, S wave velocity; p, density; Q•,

P waveattenuation' Q•, s waveattenuation. Shot elevation

= 1197 m.

?Shot elevation = 1310 m.

error bounds can be placed on our derived models. Sensitivity studies based on systematically perturbing our final velocity models show that upper crustal velocities can vary

by ---0.05km s-• withoutappreciable misfitto theobserved travel times and amplitudes, while lower crustal velocities distinguish between mechanismsbut only considered effec-

can vary by -+0.15km s-•. As a final comparison, our

tive attenuation.

When using the reflectivity method, the velocity model is approximated by homogeneousplane layers, and gradient zones are approximated by dividing the model into thin layers [Mt;iller, 1985]. Layer thicknesses were chosen so there were approximately four layers per wavelength at the highest frequency of the source. Testing of the reflectivity algorithm showed that approximating the gradient zones by finer subdivisionsshowed no change in the calculated seismograms. S wave velocities in each layer were scaled linearly to the layer P wave velocity assuminga Poisson's constant of 0.25. Following Nafe and Drake [1957], layer densities were determined using the empirical velocitydensity relationshipp = 0.252 + 0.379 a, where a and p are the layer P wave velocity and density, respectively. The seismic source was a Ricker wavelet whose frequency content was matched to the spectral content of the source. The synthetic seismogramswere calculated in units of velocity. All synthetic seismogramswere calculated over a

phasevelocitywindowof 4.5-20.0km s- • in orderto include the Pg, PiP, PmP, andPn phases.The directP wavewas excluded from the synthetic seismogram computation by selectingan apparent velocity window larger than the velocity of the surficial sediments. In the Basin and Range, where effective attenuation of high frequencies is commonly observed [Braile, 1977], it is important that the seismogramsaccurately represent anelastic attenuation. For this purpose, causal absorption was modeled in the synthetic seismogramsusing the frequencydependent, complex velocity law of O'Neill and Hill [1979]. For the synthetic seismogram simulations shown here, the

TABLE

z, km

a, kms-1

lb.

East-West

/3,kms-1

Profile

p,g cm-3

0.00

2.80

SPI* 1.62

1.31

200

100

1.20 1.20 2.50 6.00 8.00 19.00 19.00 29.00 31.00 40.00

3.60 5.60 5.85 6.05 6.10 6.25 6.40 6.70 8.00 8.02

2.08 3.23 3.38 3.49 3.52 3.61 3.70 3.87 4.62 4.65

1.62 2.37 2.37 2.54 2.56 2.62 2.68 2.79 3.28 3.30

200 200 200 200 400 400 1000 1000 1000 1000

100 100 100 100 200 200 500 500 500 500

0.00 1.60 1.60 3.00 8.00 21.00 21.00 22.00 22.00

2.80 3.50 5.60 5.95 6.00 6.20 7.25 7.22 6.40

1.62 2.02 3.23 3.44 3.46 3.58 4.19 4.17 3.70

1.31 1.58 2.37 2.51 2.53 2.60 3.00 2.99 2.68

200 200 200 200 200 200 400 400 1000

100 100 100 100 100 100 200 200 500

30.00 32.00 40.00

6.75 7.90 7.92

3.90 4.56 4.59

3.87 3.25 3.27

1000 1000 1000

500 500 500

SP7P

See Table la footnote. *Shot elevation = 1188 m.

?Shot elevation = 1932 m.

21,830

BENZ ET AL.' BASIN AND RANGE CRUSTAL STRUCTURE

NORTH-SOUTH

typical of semiconsolidated to consolidated Tertiary and Quaternary basin sedimentsfound in the Basin and Range. The crystalline upper crust, deduced from SP8 (Figure 2), is characterized by velocities increasing with depth from

PROFILE

A SP8 ,

I

.

I

,

I

.

SPll I

5.40km s-1 at 1.5km to 5.95km s-• at 5.0 km depth.The •'

rapid velocity increase is considered the result of crack closuresand the consequentialdecreasein porosity, a property common to all of the velocity-depth functions shown in Figure 6. From 5.0 to 18.5 km, velocities increase slowly to

Layer2

6.22 km s-1. For SP8, the midcrustalto lower crustal boundary is modeled as a first-order discontinuity where the

ao40

velocityabruptlyincreases from6.22 to 6.40 km s-1. This

'"

small midcrustal discontinuity produces the PiP phase seen on the SP8 record station (Figure 2). From 18.5 to 30.0 km the velocity-depth function increases linearly (with a higher

- Layer 5 ..........

3

4

5

6

7

VELOCITY

gradientthanfor theuppercrust)to 6.85km s-1 .

8

(km/s)

VELOCITY

EAST-WEST

The crust-mantle transition occurs over a depth of 4.5 km and includes two velocity discontinuities (Figure 6a). The first discontinuity is modeled as a 1.0-km-thick transition zone in which the P wave velocity increases from 6.85 to

(km/s)

PROFILE D

SP1

-•''

',•'

7.65km s-1. A 3-km-thicklayer,with an averagevelocity 7.65km s-1 is underlain by a secondsmalldiscontinuity at 34 km depth,wherethe velocityincreases to 7.9 km s-1.

SP7

.....

-•.z ' ' ' ß'''''

This rather complex crust-mantle transition, which is only seen on the north-south profile, is based on the observations

10-

of (1) a highapparentvelocity(--•7.6km s-1) wide-angle reflection 20-

observed

from 140 to 180 km from both SP8 and

SP11, and (2) a weak Pn refraction with an apparent velocity

of--•7.9km s-1 thatarrivesbeforethewide-angle phase.

--

The velocity-depth function derived for SP11 (Figure 6) is similar to that of SP8. Near SP11, the crystalline upper crust

30-

Altem

4O 3

4

5

6

7

8

VELOCITY (kin/s)

3

4

5

6

7

8

VELOCITY (kin/s)

Fig. 6. P wave velocity-depth functions determined from iterative travel time and amplitude modeling of SP8 and SP11 along the north-south profile and SP1 and SP7 along the east-west profile. As a comparison, the dashed pattern represents the extremal bounds derived from tau-p inversion of the same data [Hawman et al., 1990]. SP7 was not consideredin the study of Hawman et al. [1990].

begins at 1.6kmwitha basement velocityof 5.50kms-1 that increases to6.0 kms-1 at a depthof6.0km.From6.0to 17.0 km,thevelocityincreases to 6.25km s-1. Dueto thelackof an observable PiP phase on data from SP11, no midcrustal discontinuity is modeled, but there is a change in velocity

gradient,andthevelocityincreases to 6.85km s-1 at a depth of 28.0 km. Like the velocity model derived from SP8, the crust-mantle transition occurs over a depth range of 5 km and consistsof two velocity discontinuities and an intervening layer. The discontinuities are from boundaries with

velocitycontrasts of 6.85-7.6km s-• and 7.7-7.9 km s-• velocity-depth functions from SP1, SP8, and SPll are compared directly to the extremal bounds derived from tau-p inversion of the same data [Hawman et al., 1990].

North-South Profile

SP8 and SPl l velocity-depthfunctions. Shown in Figure 6 are the velocity-depth functions derived from iterative amplitude and travel time modeling of SP8 and SP11 on the north-south profile. The velocity and attenuation parameters used in the synthetic seismogramcalculations are also listed in Table 1. In all cases, depths are measured relative to the shot point elevation. For SP8 and SP11, elevation above sea level is 1197 and 1310 m, respectively [Whitman and Catchings, 1987]. The surficial geology along the north-south profile is relatively uniform (Figure 2) with most stations on or near bedrock along the edgesof the alluvial-filled valleys. Consequently, the surficial velocities are similar between

shotpoints.Sediment velocities rangedfrom2.8 km s-1 at the surfaceto about3.6 km s-1 at a depthof 1.5km andare

that occur at depths of 28 and 33.3 km, respectively. The dominant crustal P wave phases, seen in both the

observed andsyntheticseismograms, are the Pg andwideangle PmP phases (Figures 7 and 8). The most significant difference between the two synthetic sections is that the model for SP8 produces a distinct PiP phase, while the model for SP11 does not. The apparent velocity of the Pn phase, from both SP8 and SPll, is approximately 7.9 km

s-• , which suggests minimaldip along the north-south profile. A comparisonof the observedand synthetic seismograms in the range 120-200 km shows that the synthetic seismogramscorrectly model a large amplitude phase that

closelyfollowsthe weak Pn phase.The apparentvelocity of

this phaseis unusually high,approximately 7.6 km s-1, indicating that the phase bottoms in a high-velocity region. The simplest model that can produce this observation is a small velocity discontinuity at a depth of 34 km, as shown in Figure 6. Wide-angle reflections from this boundary will have high apparent velocities and because of the thin transition layer, the travel time will closely follow the upper mantle head wave, Pn, that is observedat a velocity of 7.9

BENZ ET AL.: BASIN AND RANGE CRUSTAL STRUCTURE

6

N SP8

TRACE NORMALIZED

ca LJ •

21,831

'

P

PiP

t

•- '0 -2 6

5



0 -2

-200

-t80

-180

-ltt0

-120

-100

-80

-60

-tt0

-20

0

D I STRNCE

Fig. 7. Comparison of SP8data(top)andreflectivity synthetic seismograms (bottom)calculated for themodelin Figure6. Eachtracehasbeenband-pass filteredbetween0.5 and12.0Hz andplottedusinga reducing velocityof 6.0 km s-1. Observeddataare time term correctedto the depthof the shotpoint. SeeFigure3 for detailsof plot.

8 8P1 I

TRACE NORMALIZED N

5 o •

5

i t!



1

0

0

_! --2



6

6

5

5

0

0

-1

-1

-2

-2 0

20

tt0

60

80

100

DISTRNCF

120

1 t10

160

180

200

(KM)

Fig. 8. Comparison of SPI1 data(top)andreflectivity synthetic seismograms (bottom)calculated for the modelin Figure 6. See Figure 3 for detailsof plot.

21,832

BENZ ET AL.' BASIN AND RANGE CRUSTALSTRUCTURE

Pg AMPLITUDE v$ DISTANCE SP11 '

u.I

'

I

'

'

I

'

'

'

I

'

'

'

I

'

'

' I' ' Observed

1

Theoretical Im ß

m

0.1

C B

m ß ß

0

m

ß m

ß

0.01

20

40

60

80

DISTANCE Model

A

0

Model

kl

.10

B

(krn) Model

I I I I

400

120

100

C

Model

I I I I mI

D

I I I I I I

600

• 20

1000

m ";;;;'" 1 / lOOO '1

40

• • , 345676

345676

Vel

(kin/s)

Vel

,

,

(kin/s)

,

• • I , 345676

Vel

,



(kin/s)

,

, , 345676

Vel

(kin/s)

Fig. 9. Observed Pg amplitudes versusdistance for SPIl compared with theoretical amplitudes for fourcrustal models(A-D). Models A, B, and C (bottom) utilize higher seismicattenuationin the upper crust to model the amplitude

distance data.In modelD, Pgamplitudes aredecreased by theeffects of a crustallow-velocity zone.Numbers within the four velocity-depthplots indicateinverseP wave attentuation(Q). Model A, with Q = 200 in the upper crust, is the prefered model.

kms-•. Thephaseobserved in thesynthetic seismograms at Black Forest, West Germany, a region of crustal extension 200 km and approximately 4.0 s reduced time is a crustal P wave multiple.

and comparablehigh heat flow.

tances of 110 km from SP8 and SP11 (Figures 7 and 8) and

simulatingupper crustal P wave attenuation (models A, B,

To investigate theobserved Pgattentuation withdistance, Q structure. The Pg phaseis clearlyidentifiedto dis- synthetic seismogramswere computed for three models

the observedtravel time of the Pg phaseconstrainsthe and C; Figure 9). The attenuation modelswere computedfor velocity to a depth of 10-12 km. Without consideringattentuation effects, however, the velocity models for SP8 and SPll do not predict the observed amplitude decay with

distance of theobserved Pephaseasseenin Figures7 and8.

the velocity model for SP11 (Figure 6). In addition, synthetic seismogramswere computed for the model of an upper crustal velocity reversal at a depth of 9.5 km (model D, Figure 9). The upper crustal low-velocity zone model was modifiedfrom the SP11 velocity-depth function by decreas-

Plausible models that can explain the observed amplitude decay involve either low anelasticbody wave attentuation or an upper crustal velocity reversal that results in downward

ingthevelocityfrom6.05km s-• at 9.5 kmto 5.95km s-1

refractionof the Pe phase,therebyreducingthe Pe ampli-

entire crust.

tude. The latter mechanism was successfully used, for example, by Gajewski and Prodehl [1987] to model the

The amplitudesfrom both the syntheticseismogramsand observed data were determined by measuring the peak-to-

at 10 km depth. Q for this model was chosenas 1000for the

overthe firstfull cycleof thePe waveform. amplitude decrease ofPefromrefraction dataacquired inthe peakamplitude

BENZ ET AL ' BASIN AND RANGE CRUSTAL STRUCTURE

21,833

Traces with a signal-to-noise ratio less than approximately observed from SP8 (north-south profile). From 19.0 to 29.0 1.5 were not included. All seismogramswere recorded on km the velocity-depth function increaseslinearly to 6.70 km instruments of the same type; therefore no instruments s-• at the baseof crust.The crust-mantle boundaryis correction was applied other than removing differences in modeled as a 2.0-km-thick transition zone, over which the gain setting. velocityincreases from6.70to 8.00km s-•. Pn is weakbut Figure9 showsa comparison oftheobserved Paamplitude observable at approximately 0.0 s reduced time at 150-km and the synthetic amplitude-decay curves calculated for offset. From a depth of 31.0-40.0 km, the mantle is modeled

modelsA, B, C, and D. The observedPa amplitudewas by a slightvelocityincreaseto 8.05 km s-•. Unlike the determinedfrom the first full cycle of the phase, as observed on the unfiltered data, and the theoretical curves were

calculated by measuring thepeak-to-peak Pa amplitude from

north-south profile, we do not recognize a crust-mantle transition with two separate velocity discontinuities. For offsetsgreater than 140 km, the wide-angle Prop phase has

anapparent velocitycloseto 6.7 km s-I , implying propagaamplitudes were normalizedrelativeto the maximumPa tion in a lower crust with a lower seismic velocity than the synthetic seismograms. The observed and theoretical

amplitude between 15 and 120 km. The best fit model is model A; both models B and C underestimatethe amplitude decay with offset, while model D over predicts the amplitude decay at large offsets.

observed on the north-south profile. Synthetic seismograms of the velocity model for SP1 (Figure 6) are compared to the observed refraction record sectionsin Figure 10. For SP1, a comparison of the observed Bandaet al. [1982]pointsoutthatPa amplitudes cannotbe and synthetic seismograms show that travel-times and amused exclusively in modeling apparent Q, because subtle plitudes are in agreement to offsets greater than 200 km. changes in upper crustal gradients can cause changes in Three significant features are seen in the observed and amplitude-distancebehaviour. Although a low Q of 200 in synthetic seismogramsfrom SP1. First, the midcrustal vetheuppercrust(modelA) bestfitstheobserved Paamplitude locity discontinuity does not produce a prominent phase decay, the synthetic seismogrammodeling showed that a Q similar to that observed along the north-south profile. Secof 200 for the entire crust producestoo much attenuation of ond, PrnP propagates with a low apparent velocity at large other crustal body wave phases,suchas the wide-anglePrnP offsets ( 100 km, and proximity in time of the PmP and PiP phases makes it difficult to distinguish them separately beyond 100 km. A mantle Pn phase is not clearly observed from SP7, suggest-

7

15

ingeithera slightmantlegradientlessthan---0.002s-1 or a

•,20

MODEL

•. 25

30

highly attenuating mantle region. The data are not sufficient to distinguish these effects. Constraints on lower crustal velocity gradients. In order to assesslower crustal gradients, we have calculated synthetic seismogramsfor three plausible end-member models, each of which fits the observed seismic travel times (Figure

2

MODEL 3 MODEL

I

12).Thesemodelsare (1) a moderategradient(0.022s-l, model1), (2) a highgradient(0.030s-l, model2), and(3) a lowgradient(0.017s-l , model3) from19to 25 km followed by a very highgradient(0.175s-l) from 25 to 29 km. All

Fig. 12. Three alternative lower crustal velocity models that fit the observed refraction data from SP1, these models are (1) a

three

models

have

a 1.5 km-thick

crust/mantle

transition

moderate lineargradient (0.022s-1, model1),(2) a highlinear zone.A modelwith a low (lessthanabout0.015s-l) does

gradient(0.03s-', model2), and(3) a variablegradient(0.017s-1 not fit the seismic travel times of the PmP (Moho) reflection between19and25 km and0.175s-1 between25 and29 km, model at wide angles.

3).

The travel time of the lower crustal reftaction (Pi, a lower crustal turning phase) and the PmP reflection for these three manner compatible with the other velocity-depth models) to models can be distinguished from each other (Figure 13). In increase linearlyfrom6.40km s-l to 6.75km s-l at a depth model 1, these two phasesmeet at a cusp at ---235 km, while of 30.0 km. The crust-mantle transition is modeled by a in model 2 this cusp is located 30 km closer, at 205 km velocityincrease to 7.90km s-l at 32.0km. Thewide-angle (Figure 13). In model 3 the low seismic gradient from 19 to 25 reflection from this transition zone is observed as the relakm results in a greater distance of the cusp to beyond 280 tively low amplitude phase beginning at 80.0 km distance km. The high-velocity gradient from 25 to 29 km does not

õ



MODEL 1

TRACE NORMALIZED

1



0 -1

-2 I

I MODEL

6

I

I

I

I

I

I

I

I

I

I

I

I

2

6 5 m•

2

x

2

i

1

1

0

0



-1

-1

-2

-2

MODEL

6

3

5

3



0

-1

-12

-2

0

20

m.t0

60

80

100

120

ll-t0

DISTRNCE

160

180

200

220

240

260

280

(KH)

Fig. 13. Reflectivity synthetic seismogramsfor the three velocity models shown in Figure 12. Upper crustal velocities are the same as that for SP1. Dots represent the cusp of the PmP reflection predicted from travel-time curves of the three velocity-depth functions shown in Figure 12, respectively.

21,836

BENZ ET AL ' BASIN AND RANGE CRUSTAL STRUCTURE

produce an amplitude or travel time cuspwithin a distanceof 280 km. Model 3 can be distinguishedfrom the other models by the lower apparent velocities obtained from the lowgradient zone between 19 and 25 km (Figure 13). These forward calculations indicate that it is very difficult

clearly observed on the north-south profile (SPll and SP8, Figure 6). It consistsof a 3-5 km thick layer with an average

velocityof 7.6 km s-i; thesedatado not providereliable estimates of the velocity gradient in layer 4. Layer 5, with a

velocityof 7.8-8.0km s-1 , is identifiedas the top of the

to resolve the structure

of the lower crust based on seismic upper mantle. The small amplitude of the Pn phasesuggests in thislayeris low, 0.002s-1 or travel times and amplitudes of P waves alone. Furthermore, thatthe velocitygradient the subtleamplitudeeffectsof differinglower crustalseismic less, in the upper mantle. Previous P wave velocity models of the Basin and Range gradients are likely to be smaller than the effects of lateral velocity variations over the distancesconsideredhere (280 and surrounding regions were based on seismic data ackm). Precise travel time measurements(i.e., reversed appar- quired in the 1960s and 1970s that had about one-fifth the ent velocity determinationsfor Pi andProP) providethe best station density of the current observations, typically 5-km constraintson lower crustal seismicgradients. In the present station spacings and 200-km shot spacings [e.g., Pakiser, 1963; Eaton, 1963; Hill and Pakiser, 1966; Prodehl, 1979; case, we have determined a seismicvelocity at the top of the lower crust of about 6.4 km s-1 that increasesto 6.6-6.8 km Pakiser, 1985]. A comparison of these results with previs-1 at a depthof 27-29km. A high-velocity transition zone ously determined crustal models provides a regional per(6.9-7.5km s-•) at thebaseof thecrustispermitted butnot spectiveon the crustal structure of the Basin and Range. The Basin and Range is bordered on the east and west by the required by the data for the east-west PASSCAL profile. A basallayerwitha velocityof about7.5kms-1 isrequired on Colorado Plateau and Sierra Nevada, respectively, geologithe north-south profile because a clear, separatephase from cal provinces with -> 40 km thick crust (profiles A and H, Figure 14). Upper crustalvelocitiesin the Sierra Nevada and this layer has been identified (see discussionabove). the Colorado BASIN AND RANGE CRUSTAL VELOCITY STRUCTURE

Some inferences

AND COMPOSITION

Plateau are similar to those of the Basin and

Range, but the upper crust is generally thicker, 21-27 km versus 15-17 km in the Basin and Range. Significantly, no

basalcrustallayer(7.3-7.6 km s-1) hasbeenreportedin

can be made from the 1986 PASSCAL

seismicexperiment regarding the compositionand physical propertiesof the crust and upper mantle in the northwestern Basin and Range, importantly, how crustal compositioncan be determined from seismic velocity models and how the seismicvelocity structure can be related to the processesof lithospheric extension are considered. In this section, we will (1) summarize the essential features of the velocity structure and compare it to crustal models from the surrounding region, (2) present some inferencesregardingthe composition and physical properties of the crust and upper mantle, and (3) discuss implications regarding Basin and Range crustal extension. The crustal and upper mantle structure that has been obtained here can be divided into five layers (excluding the surficial sedimentary layer), each with a positive velocity gradient (Figure 6a). Layer 1 lies directly beneath the surficial unconsolidated sediments and likely consists of fractured basement rocks. This layer has a high-velocity

these two bordering geologic provinces. However, such a layer, where thin (2-5 km), is difficult to detect in areas of thick crust and especially with sparse data coverage of the earlier seismic profiles. The crustal velocity structure derived from the 1986 PASSCAL Basin and Range seismicexperiment is similar in several important aspects to previously determined crustal models for the Basin and Range of Nevada and Utah (profiles B, E, F, and G, Figure 14). The seismicrefraction profile nearestto the PASSCAL investigationsis the FallonEureka profile (profile B, Figure 1), interpreted by Eaton [1963] and Prodehl [1979]. Most recently, Catchings [1987] interpreted this profile and suggested a laterally varying crust that can be summarized by two crustal columns representingthe west (Fallon) and east (Eureka) ends of the profile. These models show an approximately 19-km-thick

uppercrust(6.25km s-l), a 7-km-thicklowercrust(6.8km s-l), anda 4 to 10-km-thick basalcrustallayer,7.5km s-1 (Figure 14, profile B) [Catchings, 1987]. Thus, the average

gradientof 0.12 s-I andan averagevelocityof 5.7 km s-• Fallon-Eureka model (Figure 14, profile B) is nearly identical (5.45-5.95km s-l). Layer2, at 5-18kmdepth,togetherwith to the SP11 velocity model (Figure 14, profile D), which it layer 1 comprise the upper crust. The average velocity in

crosses,but differs from the SP1 model (Figure 14, profile C)

of the 7.5 km s-I layer. layer2 is 6.1 km s-I (5.95-6.25km s-l). Layers1 and2 are in thethickness A comparison of our northwestern Nevada generalized crustal columns(Figure 14, profiles C and D) showsthat the 0.12s-l in layer1 to 0.02s-1 in layer2. Layer2, likelayer thickness and velocities of the upper and lower crust are 1, can reasonably be inferred to consist of basement rocks similar to that observed for east-central Nevada (Figure 14, with most cracks and fractures closed. profile E), as derived from a refraction profile recorded Layer 3, the lower crustal layer, at 18-29 km depth, has an between Eureka, Nevada, and Boise, Idaho [Hill and Paaverage thickness of 11 km, an average seismic velocity of kiser, 1966]. No transitional lower crust-upper mantle is 6.6 km S-1 anda velocitygradientof 0.04s-I Layer3 is observedfrom the Eureka-Boise profile, but such a transidistinguished from layer 2 by (1) either a small velocity tion cannot be ruled out because of the wide station spacing discontinuity (about0.2 km s-1) or, whereno boundary used to record this profile. The crustal velocity structure of the eastern Basin and occurs, by a higher seismic velocity gradient; and (2) an averageseismicvelocitythat is 0.5 km s-• higherthanin Range is known primarily from unreversed profiles and is layer 2. Likely possibilities of lower crustal composition henceequivocal (Figure 14, profiles F and G). Neither two of (layer 3) are discussed below. the eastern Basin and Range profiles were recorded wholly Layer 4 is a crust/mantle transition layer and is most within the Basin and Range (profiles Ga and Gb, Figure 14) not separatedby a seismicboundary, rather they are distinguished by a reduction in the seismic velocity gradient from



ß

BENZ ET AL.' BASIN AND RANGE CRUSTAL STRUCTURE WESTERN

A

B

Sierra Nevada Wesleto Nevada

I•

U.S. CRUSTAL

C

D

NW Nevada

NW Nevada

This Study

This Study

, b_

sP._2L sP,,

21,837

STRUCTURE

E Eas!-Cenlral

F Wesleto U!ah

G

H

Cenlral Ulah Colorado Plaleau

Nevada



, b

i:." ?.%:.".;'..'¾' •.ll • i'•.......:.,....: :.; • ....•'.;" •? •'.•.'..:;"..• i

/ l;•i:•/i•:!i:!;:i:i"-.:i:i t•i••

?,:,., ;•

•:'""•

:::::::::•':"::';'::";:::

. .. '•;;'::.-'•..::!i:i! ,o1I::.•!.....:•i!:-• I.:"•.......':.....'••!! '•..'•;i::! ':'""'" '•-•:•;: '•'",•::::! '-' :":: •'"" '":•': •

:"'•""'"' ":' '•,.,, ":"'"-"' "-' "' ".' )-::i!: ,o

!::!::";•:?'i:i

/ [;:-.-:-.-::

•' ,o-IE:i:!::-:-: t;,:•' •..: :: ,:-:-• !: ,o ,,. '-'-',-:-, ,. ".:'.."..•; -.,• .......... .. .... .-.-.v.-:-:.:: . :-,:•'.': •';,.,'..".' ..............••::::i'•'•'•'"'•""""":•'--'•!

•,

.....

:::::::::: ::',., :: ,,,':': :i:::::::i .......... .......... ,...-...:.: :.•.?:. .:•.•s-: ::::::::.:. .-.-,;,;;. x •;•;i\ ? u•3o .4[:•:•:!:•:• / IZ..-s., '2;--;,,-;•--i•==•==;•,•,---?:-:-:-:-:-: ?''liii'iiiiil-"

'-

I I::::'.,'.-:i:i/ P.,'., :v I

t:-.,'-.•:.,::: ß;,;-;-;,;

/

fl s o s o

30

?.9

ß .

50- :'i•'9"• IIII

-50

• 3.0-5.9:/.•i• 5.9-6.3!• 6.3-6.8•r• 6.8.7.2•;• 7.2.7.6[• >7.6 Supracrustal, fractured basement, and

crustal

Felsic

Intermediate

Mafic

MaficlUltramafic

Ultramafic

low

velocity zone (LVZ)

Fig. 14. Generalized crustal velocity columns for the Basin and Range province (B thru G), Sierra Nevada (A), and Colorado Plateau (H). See Figure 1 for references.

seismic velocityof 6.1 km s and one of the profiles never attained a Pn velocity greater Table1) andhasanaverage than7.4 km s-1 . The Delta-Westprofile(Figure14,profile The upper crust of the Basin and Range is similar in velocity F) was recorded westward across the area of greatest late Tertiary extension but was not reversed and has large station spacings.Despite these limitations, some comparisonsare possible. The upper crust in central Utah has a thickness comparable to northern Nevada (15-17 km), but the three Utah crustal models show evidence for upper crustal lowvelocity zones. The seismic velocity of the lower crust

to the Colorado Plateau but is 5-10 km thinner, probably as a result of stretching or attenuation by extensional faulting. 3. The middle crust is 7-14 km thick and has an average

arguesthatthis7.4 km s-1 layeris a downdiprefraction in theuppermantlewith a truevelocityof about7.9 km s

regionof activdextension.Priorto discussing compositional

seismicvelocityof about6.6 km s-1. The middlecrustis 6-14 km thinner than the upper crust.

4. A crust/mantle transitionlayer (7.3-7.6 km s-1) is

observed from SP8 and SP11 (Figure 14, profile C) and the (6.4-6.9km s-1) agreeswellwiththatobtained fromnorth- Eureka-Fallon profile (Figure 14, profile Ba) but is either thin westernNevada(6.25-6.86km s-i). (900øC). Based on these observations, the most reliable measurements, which use three-dimensional upper crust appears to best fit a general suite of granitic- models to constrain lateral current flow, are from western dioriticto quartz-richgranuliticrocks(5.5-6.2 km Utah where Wannamaker [1983] demonstrated a normal

BENZ ET AL.' BASIN AND RANGE CRUSTAL STRUCTURE

lower crustal resistivity of > 100 ohm m, more than that required by the Hyndman and Klemperer [1989] hypothesis. In the western Basin and Range, deep geomagnetic soundings were taken more than a decade ago by Schmucker [1970] and Stanley et al. [1976]. Stanley et al.'s [1976] evaluation of the resistivity structure of the Carson Sink, although not interpreted by three-dimensional methods, does suggestlower crustal resistivities of the order of those required by Hyndman and Klemperer's [1989] model. Thus, if our lower crustal velocitiesof 6.4-6.8 km s-1 are corrected for possible high pore pressure effects, they could possibly

21,839

•-60mW m-2. Inferences on mantlecomposition, basedon measuredPn velocities beneath northwesternNevada, thus

show that the upper mantle cannot be considered anomalous but has a composition typical of surrounding regions. Extensional processes. Several workers have speculated on the relationship between the thin crust, the low upper mantle velocities, and the processesof crustal extension in the Basin and Range. On the basis of geologic studies, Thompsonand Burke [1974] and Wright [1976] suggest10% extension across the northern Basin and Range, while Proffett [1977] showed locally concentrated extension of up beaslargeas6.7-7.1kms-1. Depending onhowlowcrustal to 35% in the vicinity of northwestern Nevada. Coney and pore pressure might vary with depth, lower crustal gradients Harms [1984] calculated 40% province-wide extension based could vary between the end-member velocities above. Be- on geometric models of postfault separation and fault-block cause we know little about how pore pressure varies with tilt. Wernicke and Burchfiel [ 1982] also show extension of up depth and have few constraints on inferred lower crustal to 40% on the basis of normal fault offset and tilt of adjacent pore pressure from electrical measurementsin northwestern blocks. These values are compared to inferences of regional Nevada, it is difficult to speculate beyond our simple pore extension of 10-20% from heat flow models [Lachenbruch and Sass, 1978] and •- 10% by conversion of seismic moment pressure model to explain the high-velocity gradients. Petrologic models of the Basin and Range generally invoke rates of historic earthquakes to regional deformation rates the intrusion of magmas, both within and underplating the [Eddington et al., 1987]. In comparison, the eastern Basin crust [Lachenbruch and Sass, 1978]. The high-velocity gra- and Range has experienced greater extension, in excess of dients in the lower crust of our model may reflect increased 100%, in central and western Utah [Allmendinger et al., 1983; Planke and Smith, 1990]. mafic contributions. Mafic dike intrusions throughout the Our crustal velocity model (Figure 15) shows several entire crust would significantly increase the average velocity compared to the normal continental values. Our upper important properties that bear on models of extension in the crustalvelocities, 6.2-6.3km s-1 at 15-20km depth,arenot Basin and Range. First, the crust is thin, ---30 km, compared compatible with significantamounts of mafic intrusive rocks, to 40 km or greater beneath the Sierra Nevada and Colorado and excluding possible porosity effect, our lower crustal Plateau. Assuming an ---40 km paleocrust [Coney and velocities seem rather low for major contributions of mafic Harms, 1984] requires an •-25% reduction in thickness and rocks. We can not say how much this process has affected concomitant east-west 25% extension. A higher percentage the crustal structure because deep crustal refraction and of extension is possible if crustal thickening due to magmatic reflection data are not sufficientto unambiguously reveal the intrusion accompanied extension. Second, our velocity distribution and scale lengths of lower crustal heterogene- structure showslarge positive velocity gradients in the upper ities produced by mafic intrusions.Lower crustal reflectivity and lower crust but without major first-order velocity disobserved in the coincident reflection [Jarchow et al., 1990; continuities (exceptfor thesmall0.2km s-1 increase locally Hawman et al., 1990] data and COCORP data [Klemperer et in the midcrust, Figure 6). The positive high-velocity gradial., 1986] does suggest some component of mafic lower ents throughout the crust argues that the composition crustal intrusion, but the amount is undetermined. changes systematically with depth. This systematic change Given the wide range of velocities for a given rock type or is likely the product of continued extension and intrusion metamorphic grade, it seems that the first possibility, a into the lower crust. In addition, high effective pore pressure systematic variation in composition with depth, is the most in the lower crust, due to progressive metamorphism in likely explanation for the modeled seismic velocity gradient. lower crust, may also contribute to the high seismic gradiIn this model, the lower crust would grade with depth from ents. We interpret the Moho in northwestern Nevada to be a an intermediate composition to a mafic composition. The mechanism would be the intrusion of basaltic melts into the transitional zone, 1-5 km thick, beginning at depths of •-30 lower crust. The correspondingtransition zone, the Moho, is km. The Moho, as defined by Hague et al. [1987] from the likely composed of mafic and ultramafic rocks (6.8-7.6 km deep seismic reflection data, is placed at the bottom of the s-i), whilethe uppermantlemustbe ultramafic in compo- highly reflective lower crust. The correspondence of Hague sition. et al.'s [1987] Moho is generally with the top of our transiThe observedPn velocity in northwesternNevada may be tional Moho layer. Klemperer et al. [1986] identified a band corrected for temperature by calculating the differential of discontinuous reflections ---5 km above the reflection

temperature betweena 90 mW m-2 geotherm for the Basin andRangeanda 60 mW m-2 continental averagegeotherm

Moho as the "X"

reflection.

We see some weak evidence

this zone on the basis of discontinuous

reflections

about

for 1 s

ahead of the PmP reflector at 80-100 km on SP11 (Figure 3). However, this phase is not correlatable over most of our record sections, does not appear to be a prominent phase, 10-4 km s-1 øC-1 [Christensen, 1979].Our temperature-and is not readily modeled deterministically. correctedPn velocitiesfor northwesternNevada range from 8.04 to 8.20 km s-1. These are values close to the North

[Lachenbruch and Sass, 1978], at a depth of •-30-34 km from which the Moho is first observed, and by correcting the velocity assuming a mean temperature coefficient of 5.5 x

American meanPnvelocityof 8.02-+0.21km s-1 [Mooney and Braile, 1989] and similar to Pn velocities found in the midcontinental region, where the mean surficial heat flow is

CONCLUSIONS

The crustal velocity structure of northwestern Nevada is relatively homogeneousand is distinguishedby a positive

21,840

BENZ ET AL.' BASIN AND RANGE CRUSTALSTRUCTURE

crustalvelocitygradientrangingfrom 0.02 to 0.04 s-1 . Crustal thickness averages ---30 km and varies by less than 5 km over the 280 km east-west profile length and 3 km over the 220 km north-south profile length. Crystalline upper crustal velocities (

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