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• #### Supplementary: Plate 11-1 198-198

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III # INTRAPLATE TECTONICS

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Tectonics of Noncollisional Regimes The Modern Ancles and the Mesozoic Cordilleran Orogen of the Western United States 6 INTRODUCTION B. CLARIS BURCHFIEL Massachusetts Institute of Technology An early hypothesis of plate tectonics was that plates moved as rigid pieces of lithosphere and that the relative motion between plates was talcen up at narrow zones along their boundaries. Later studies suggested that inter- action at plate boundaries could produce deformation, magmatic activity, and metamorphism for a considerable distance from those boundaries. Simply stated, the prob- lems to be addressed now are "what intraplate features and events are the result of plate-boundary interactions, what are the interrelations between these features and events, and what are the processes that cause them." All three types of plate boundaries, transform, subduction, and spreading, show widely distributed intraplate effects. In some cases the boundary can be defined along a narrow zone, but in others it is a broad diffuse zone of plate interaction. This latter type of diffuse or soR boundary is common in which one or both of the plates is composed of continental lithosphere. The interrelations between all the geological features and events at these diffuse bound- aries are unclear. The study of intraplate or dif~se-plate 65 boundary features and events within continental crust should be an important focus for any program in crustal dynamics. It is in this setting that most continental crust is formed and modified. The subduction of oceanic lithosphere is considered a stable process because it is denser than the astheno- sphere; thus, subduction could extend over long time periods until it is terminated by changes in relative plate motion or by collisional processes. It is evident from studies of modem plate boundaries that subduction of oceanic lithosphere can cause a variety of structural ex- pressions in an ovemding plate. Extension in He over- riding plate is expressed by the formation of back-arc basins and marginal seas. Oblique subduction can pro- duce strike-slip faulting. Horizontal compression in the overriding plate can lead to the development of complex erogenic belts. All three of these types oftectonic activity in an overriding plate can be coupled in venous ways to produce a wide range of complex Reformational styles. Even though the geological and geophysical evidence in- dicates that these types of structural effects exist in over-

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66 riding plates, the condition that led to different tectonic settings is effectively unknown. This chapter examines the nature and extent of plate-boundary-related effects within continental lithosphere. The focus will be princi- pally on the setting where oceanic lithosphere is being subducted beneath continental lithosphere and the leading edge of the continental lithosphere is under compression. To demonstrate the presence of compression in the overriding plate, two lines of evidence can be examined: (1) modem intraplate seismicity and (2) geological inter- pretation of erogenic belts. From the seismic evidence it is reasonably clear that an overriding plate can be under compression. The interpretation of the geological evi- dence is much more complex and must be examined in detail. Interrelations between all the elements of an oro- genic belt and their relation to the plate boundary must be established before a conclusion concerning compression within Me continental lithosphere can be established. SEISMIC EVIDENCE FOR INTRAPLATE COMPRESSION o Parts of We Andes are a modern example of a noncollision erogenic belt. Stauder (197S), Dom a study of seismicity in the Peruvian Andes, has conclucled that the Nazca plate is thrusting under the South American plate beneath central and noncom Pem along a surface of shallow (1~15°) dip. Of particular importance to an understanding of Andean orogenesis is Stauder's additional conclusion that the leading (western) edge of the South American plate is uncler east-west horizontal compression for a distance of approximately 700 km east of the Pent Trench (Figure 6.1~. Numerous hypocenters are present in that plate at depths of 1 90 km and for distances up to 700 km east of the trench. Most lie at distances from the trench of 2~0 km, depending on latitude. Focal mechanisms for nine of ten of these intraplate earthquakes yield hori- zontal compressive stress axes oriented approximately east-west. The tenth shows normal faulting and is near the trench. Of the nine solutions that show horizontal compressive stress axes, three solutions indicate predomi- nantly strike-slip faulting and six indicate reverse Slip faulting. The Andes, a chain characterized by voluminous mag- matic activity, are bounded on the west by the Peru~hile trench, a zone of ongoing plate convergence, and on the east by an active fold and thrust belt, which separates orogen from craton. Rates of subduction are on the order of 7-12 cm per year and can reasonably be extrapolated back at least into the Late Miocene on the basis of marine magnetic anomaly studies. Since Middle Miocene time the central Andes, including Peru, have experienced in- tense igneous activity and accompanying orogenesis. Folds, reverse faults, and thrust faults win a relative east- ward sense of motion have developed along the eastern margin of the Andes in the sub-Andean zone (Audeband B. CLARK BURCHFIEL et al., 1973~. The Andes were uplifted to their present height during this late Cenozoic episode of orogenesis. Thus one interpretation is that the Andean orogenic belt is the result of defonnation related to horizontal compres- sion in a noncollisional subduction system. S~uder's (1975) data are preliminary, and studies ofthis type need to be considerably expanded to establish the state of stress within the leading edge of the plate and its relation to areas of modem magmatism and deformation. Even though horizontal compression can be established, how Me stress field is generated is unknown. Studies directed.toward understanding the origin of the stress fields within the continental crust at convergent bound- aries should be encouraged. GEOLOGICAL EVIDENCE FOR INTRAPLATE DEFORMATION RELATED TO NONCOLLISIONAL PLATE CONVERGENCE lithe geological evidence for compression with an over- riding continental plate is complex and less clear than the seismic evidence. The Mesozoic Cordilleran orogen~c belt of the western United States is an older and more deeply eroded Andean-type mountain belt (Hamilton, 1969; Burchfiel and Davis, 1972, 1975~. South of the latitude of central Oregon, the geological history of the Cordilleran erogenic belt suggests that the eastward sub- cluction of an oceanic plate beneath the North American plate occurrecl contemporaneously with defonnation and magmatic activity that at times extender! more than 1000 km into the North American plate. Evidence suggests that major collisional events were rare or nonexistent along this part of the plate boundary during Mesozoic time. The Mesozoic Cordilleran erogenic belt can be divided into four terranes: (~> ~ western terrane of accreted oceanic rocks, (2) a central pragmatic arc or superposed arcs, (3) an eastern terrane of east~irected thrust faults and related structures, and (4) a locally developed terrane lying east of the thrust belt of Trusted Precambrian base- ment rocks in the Colorad - Wyoming area (Figure 6.2~. The first three terranes shiR spatially with time, but events in each terrane have a crumple contemporaneity (Figure 6.2~. The fourth terrane is restricted in time, latest Cretaceous to earliest Tertiary, to a period when the Corclilleran erogenic belt underwent significant but poorly understood changes. Even Cough Me geological features and events in these four belts are broadly con- temporaneous, their interrelations and their relation to plate-boundary interaction are not well established' par- ticularly for the terranes farther removed from the plate boundary. In earliest Mesozoic time, the western boundary of the North American plate lay along a line from central Oregon through the central Klamath Mountains, central Sierra Nevada and west ofthe San Gabriel Mountains into north- eastern Baja California (Figure 6.2~. The plate boundary is at present very sinuous, which is a result of post-

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Tectonics of Noncollisional Regimes . - _> o Hi:\ J , 10° \ 'I ·--= BOL/ V/~ \\ ~ 500 KM \ 7so B ~ too 300 soo 7~M B o - . . ,, at, .. .;.; . . -, ; j o . ..... . r i . ~ ~ At. 200 1~ . , , , ~ 200 FIGURE 6.1 NIajor structural elements of the Peruvian Andes. The belts of Miocene to recent magmatic activity and deforma- tion in Me sub-Andean zone are within Me South Amencan plate and lie east ofthe Peru trench, which marks the site of subduction of the Nazca plate. Section B-B' incorporates selected hypo- centers within the outlined region (from Barazangi and Isacks, 1976). T shows the location of the Bench. Hypocenters define a shallow dipping subduction zone and distribution of intraplate seismicity within the South American plate. Mesozoic deformation. Removing later defonnation would straighten the boundary considerably. Nearly all the rocks west of this line are Mesozoic in age and repre- sent rocks of oceanic origin that were accreted by subduc- tion processes and partially or wholly reworked to form either transitional or continental crust. WESTERN TERRANE Geological data from the western terrane indicate that accretion of oceanic rocks to the North American plate began in Triassic time (Davis et al., 1978~. Triassic cherts locally associated with late Paleozoic and possible Triassic ophiolites were tectonically disrupted and em- placed by subduction processes probably during Middle Triassic to Early Jurassic time (Figure 6.31. Locally these rocks are associated with blueschist metamorphic mineral assemblages, which were formed ~0~210 million years (m.y.) ago. In parts of these accreted sequences, exotic 67 Permian "Tethyan" fusulinid faunas are present in some limestones that are mixed with the Triassic and older rocks suggesting that far-traveled oceanic rocks were in- corporated into the accretionary wedge and reworked by later deforrnational events. Jurassic ophiolites, associated sedimentary rocks, and volcanic rocks lie above and to the west of the rocks accreted in earlier Mesozoic time and represent new additions from oceanic lithosphere to the North American plate. Mast of the Triassic and Jurassic rocks can be interpreted as accreted by eastward subduc- tion of oceanic lithosphere beneath the North American plate. Some workers have argued that one or more arc collisions may have occurred during Jurassic, particularly in Late Jurassic time (Schweickert and Cowen, 1975~. The evidence is equivocal that arc-continent collisions took place at this time, and considerably more work is needed to clarify this question. Even if an arc collision did occur in the Late Jurassic, eastward under~rusting of oceanic lithosphere dominated Triassic and Jurassic plate- boundary activity. Cretaceous and early Tertiary rocks fonn the western part of the western terrane and were accreted to the North American plate during east-directed subduction of Cretaceous and early Tertiary time (figure 6.4~. Like all accreted terranes, it is uncertain whether subduction was continuous or episodic. Blueschists of Cretaceous age are common within these rocks. Geologi- cal data demonstrate that rocks of the western terrane became youngertowarcl the west, indicating retrogression of the subduction boundary during Mesozoic time Davis et al., 19781. CENTRAL OR MAGMATIC-ARC TERRANE The central terrane consists of a magrnatic arc or super- posed arcs of Mesozoic age. Arc plutons intrude Pre- cambrian crystalline rocks in the south and Paleozoic arc rocks belonging to an arc accreted in latest Paleozoic to earliest Triassic time in the norm. All host rocks for the plutons were part of the North American plate; thus the arc was built on the western edge of the Norm American plate, and its structural sewing was similar to the modern Andes. The Mesozoic volcanic-plutonic arc began to develop in Early Triassic time, and the oldest platonic rocks date to approximately 23~240 m.y. ago (Figure 6.3~. Igneous activity occurred throughout most of.\Iesozoic time, but the degree of continuity of activity is controver- sial. Several workers have discussed this problem (Lanphere and Reed, 1973; Armstrong and Suppe, 1973), and the data suggest at least three intrusive epochs: (1) 7~106 m.y. ago; (2) 132-158 m.y. ago; and (3) an unde- fined epoch older than 160 m.y. ago, with the oldest dates at 230 240 m.y. ago. Because a few concordant age pairs fall in the intervals between those epochs, the alternative hypothesis of continuous magrnatism cannot be elimi- nated. Younger igneous rocks are also present in the west- ern United States, but an important change in magmatic and structural events occurred at about 75 m.y. ago as discussed below. Magmatic activity in the arc terrane is J

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68 FIGURE 6.2 The four Mesozoic to early Tertiary terraces of the Cordille- ran erogenic belt of the western United States. The western terrane consists of rocks acereted to the North American plate dunog Mesozoic time, the central tenant is foxed by rocks of one or more magmauc arcs, and the eastern terrane consists of an east~irected fold and thrust belt. The Colorado Wyoming Rocky Mountain terrane developed in latest Cretaceous~arly Tertiary time during important rear- rangement of Cordilleran tectonic ele- ments. The line meriting the eastern edge of the western terrane is the alp proximate western edge of the North Amencan plate at the beginning of the Mesozoic. K, Klamad, Mountains; SN, Sierra Nevada; SO, San Gabriel Moun- tains. No attempt has been made to re- move intraplate defonnation such as in the Basin and Range province except for reversing movement on the San An- dreas and some associated faults. broadly over the same time span as subduction activity in the western terrane, which has led many workers to couple the two terranes into an Andean-type arc trench system (Hamilton, 1969; Burchf~el and Davis, 1972~. EASTERN TERRANE The eastern terrane lies east ofthe magmatic arc and con- sists of relatively east-directed thrust faults and associated folds. In early Mesozoic time in western Nevada thrust faults developed within Paleozoic "eugeosynclinal" rocks and early Mesozoic back arc sedimentary and volcanic rocks (Figure 6.3~. The age of these thrust faults is Early and Middle Jurassic, and post-Middle Jurassic and pre- Middle Cretaceous. Recent work has demonstrated the presence of early Mesozoic thrusting and folding in nor~- eastern Nevada, which could range in age from Middle Triassic to Earlylurassic. In the miogeosyncline on either side of the Califom~a-south~entral Nevada state line are several large thrust faults and associated folds that involve rocks as young as Early and Middle Triassic and are cut by plutons 185 m.y. old. These thrusts belong to an early Mesozoic period of thrusting that may be earlier than or synchronous with deformation in western and north- eastern Nevada. Farther southeast in southeastern Cali- B. CLARK BU RCHFIEL ! ~ i -.-.2 -.' \) ~ ~ ~ ,. . ,, 1~ ~ ~ of . , - - , - ~ .,~7 - - - t~=,1---- - - ' - - ' - - ' ' ~ 2 - .':',' ~~ /: LIZ,... ': ~ O .. -' \- ,W'-- 'at; , , _ , ~ ' l~ J - - . . ~ . ,= - . . . , - - ., . ~ ;2.,: ,-,;, 2-' /; ~ t.~: ~ ~~ _ ~ r .. . . . . . · ~ 1~ _ ~ _ · · · — ~ - - . ,q, - - - - - - · · ~ W - - ~ - . . A=> S G ,~ - ~ ,; .. In- - - . ~ ~ ~ - , , _ o 3^^~ at. - ~ . . ~ . . .^ ;v '2 t-~7 J fornia' thrusts are cut by plutons 200 m.y. old and a newly discovered thrust that is unconformably overlapped by the Upper Triassic (?,~Lower Jurassic Aztec Sandstone. Thrust faults in southeastern Califomia involve Paleozoic miogeosynclinal and cratonal facie s and their Precam- brian crystalline basement. Recent data from southern California have demonstrated that some Reformational events are older than 23~240 m.y. Whether these various events belong to one or more episodes of deformation is not yet known. At present we group them in an undiffer- entiated period of early Mesozoic deformation, which ranges from twiddle Triassic to Middle Jurassic. During late Mesozoic time, arc magmatism encroached eastward into the region of early Mesozoic deformation, with the thrust and fold belt involving rocks even further east (Figure 6.4), except in southeastern Califomia where early and late Mesozoic deformation is superposed. The thrust faults developed from west to east until rock units transitional between miogeosyncline and craton were in- volved in thrusting. In southeastern Califomia, thrust faults strike south and southeast, leaving the Paleozoic geosynclinal terane and cutting through the craton. Nearly all late Mesozoic thrust faults in this region in- volve Precambrian crystalline rocks and strike parallel to and along the eastern edge of the late Mesozoic magrnatic

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Tectonics of Noncollis~onal Regimes Aft:: _'— , ~ . . ~_~ . =_.... . . . . . ~.~ . ~ . . . , ~ ... ~ . . . . c'- , . . ... . . . ..... ~ . ~ ^. 'I _\ \. ~ ' ~ if. I, . ~ . . . . I:-: ~ . An.- , . 2 >,-, -__ ~ ~ ~ ~,:': ll -rat ~ ~ ~ . . . . . . . . . \ . . . O IN FIGURE 6.3 Spatial distribution of the western, central, and eastern terranes from Early Triassic to Late Jurassic time. The western and eastern terranes were probably continuous on both sides of the central terrane, but only their present outcrop distri- bution is shown. arc. This major change in structural trend and style occurs near the Califomia-Nevada state line and presumably continues into Mexico, although data to the southeast are scanty. Late Mesozoic plutonism occulted farther eastward in Nevada and to the north than in early Mesozoic time, but in southeastern California and to the south early and late Mesozoic igneous activity is superposed. Associated with the Mesozoic plutonic and volcanic rocks, but character- istically lying east of their major areas of development, thus transitional between the central and eastern terraces, are numerous isolated areas of metamorphic rocks that probably represent exposed culminations of ~ metamor- phic belt that may be continuous at depth. Metamorphism of amphibolite grade affects miogeosynclinal and cratonal rocks of Precambrian and Paleozoic age as well as Meso- zoic and early Tertiary (?) granitic, volcanic, and sedimen- tary rocks. Meager age data suggest Mat metamorphism began at least 180 m.y. ago and continued into Me Ter- tiary, but it is not known if this represents one long period of metamorphism or several spatially and temporally dis- 69 tinct periods. In some areas of southeastern California and adjacent Arizona, metamorphism is mid-Cenozoic in age and is associated with extensional tectonics, not thrust faulting (see Chapters 8 and 9~. It is likely that metamor- phism is synchronous with magmatism, spans a long period of time, and will be tied ultimately to magmatic epochs. The eastern or frontal thrust belt ofthe Cordilleran oro- gen lies to the east of Me Mesozoic metamorphic terranes mentioned above. In this belt, east-directed thrust plates of Late Jurassic to Late Cretaceous age define a zone of crustal dislocation that extends from British Columbia and Alberta, Canada, to southeastern California and prob- ably into Sonora, Mexico. Dunng the Late Cretaceous an important change took place in the location of deforma- tion by Trust faulting. Formation of folds and thrusts in the eastern thrust belt ceased before the end of Me Creta- ceous in a sector from central Utah to southeastern Cali- fornia (Armstrong, 1968~. North and south of this sector of the thrust belt, deformation continued through the Late Cretaceous into Me early Tertiary and ceased in Middle 1 _— . - ~ ! = _ . ~~ : ~ '. ~.^ .- -A ~ )7l C" c, ~- Q: - ~ ~ ~! . · . . ~ ~ · · · · ~ ... ~ . ,_, . + I, ~ . . . - . . · · · · · ,, - ~ — 4 · . _ · _ . _ _ _ _ j" ~ _ - . _ _ - — — — — — - / t — - - - - · _ _ , - ~] , . ~ . — , — — — — ~~ r _ . . ~ _ . _ _ . ~ . . _ _ . ,, _ ~ _ — — ~ / , — , _ . _ , · - — ~ _ , , .— _ . . . . _ ~ r . ~ . . _ _ _ ~ I ,,,, ~ ,~ Id_ '': ' ::,:,. ~ — . _ , , , ~ ~ ~ — · _N ,: - ' _ · ~ ~ — - — - _ _ _ ~x _- · - - - - - - , Car 'W-:-:-:-:-:-:-:.-:- ~g _ _ at- . ... ~ , . ~ --; . . . - 3 0 0 ~ M . _ _ . . . _ , , . . _ . _ _ . _ FIGURE 6.4 Spatial distribution of the western, central, and eastern terranes from Late Jurassic to latest Cretaceous time. The westem and eastern terranes were probably continuous on both sides of the central terrane, but only their present outcrop distri- bution is shown.

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70 or Late Eocene time. The period from latest Cretaceous to Eocene (75~;0 m.y. ago) is Me time of He classic Cyanide orogeny. Low-angle thrust faults developed during Lara- mide deformation have a structural style similar to Hose developed in He earlier Cretaceous events but generally lie somewhat farther east. COLORADO—WYOMING ROCKY MOUNTAIN TE RBANE Lee fourd1 te'Tane is only locally developed both spatially ant] temporally ant! consists of large uplifts of Precarn- bnan crystalline rocks that extent] Tom southern Montana into New Mexico. These uplifts, commonly referred to as He Colorad - Wyoming Rocky Mountains, developed only dunng the Laramide Orogeny of latest Cretaceous and early Tertiary time. The structural geometry of He Faults bounding the uplifts has been controversial, one group postulating Trust faults Blat steepen with depth, a second group postulating that the boundary faults are moderate to low angle (for a review see Sales, 1968; Steams, 1971~. Recent seismic reflection lines across one of the uplifts, the wind River Range, by the Consortium on Continental Reflection Profiling (COCORP), has demon- s~atec! a gently dipping thrust fault bounds the uplift (see FIGURE 6.5 Spatial distribution of die western, central, eastern, and Colorad - Wyoming Rocky Mountain terranes Tom latest Cretaceous to early Tertiary time. The western terrane was probably more extensive than its pres- ent outcrop distribution and may have been continuous along the entire west coast, but it is now largely in the off- shore region. B. Cal ARK BURCHFIEL Chapter 10~. Some of the uplifts can be interpreted to be bounded by thrust faults, but several are still equivocal. The geometry of these uplifts is critical because their Origin has been ascribed to horizontal compression or ver- tical displacement without horizontal crustal shortening. The COCORP data suggest Hat the former interpretation may be correct, in which case it may be possible to relate He structural origin of this terrane to plate-boundary in- teraction even though He boundary is more than 1000 km to the west. Igneous activity In the western part of He Now Ameri- can plate also underwent a significant change during the Late Cretaceous (Figure 6.5~. Dunng most of Mesozoic time, a plutoni~volcanic Andean-type arc was active as discussed above. Although He problem ofthe continuous or episodic character of magmatism has not yet been re- solved satisfactorily, it appears that during most of Meso- zoic time magmatism was characteristic of much of He western part of the Norm American plate in the United States and southern Canada. About 75 m.y. ago a change occurred in the patting of igneous activity. Igneous rocks intruded during Amide time (75 50 m.y. ago) are pres- ent in Canada and Idaho and in southeastern California and Arizona (and presumably western Mexico), but Hey are very rare in central and northern California, Nevada, :~-~l -~# ~~\' 5~. ~ - rl1 ~ ) AN ~ - ~ \ % l \< \ o - D D .. \ \ l I I. ~~ Zip'' .. an of, ~ : cn , . of ~ ~ x, :.,'rNc.~4 Z ' ~ · 'A- ' ~ ~ . ~ . ~ . ~ . ~ ~ ~ · . . . i- , · 4, . . ~ . _ ~ ~ ~ w ; , - ~ — ~ ~ !~ ~ rn , '. D i Z rn '. .i 1\ <; v C, ~ ~ rr1 rn

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7 I NTRODUCTI ON Models for Midcontinent Tectonism WILLIAM J. HINZE and LAWRENCE W. BRAISE Purdue University G. RANDY KELLER University of Texas, E! Paso EDWARD G. LIDIAK University of Pittsburgh The midcontinent region of the United States has long been regarded as part of the stable craton. Geological evidence has led to He assumption that this area has un- dergone only minor tectonism during the past several hundred million years and that this tectonism has largely taken the fonn of broad, slow, vertical movements. However' during the past decade there has been accumu- lating geological evidence and increasing awareness that the midcontinent region has been and is at present tec- tonically active. This change in geological thought has come about because of studies of earthquake activity and enhanced discrimination of lateral crustal variations by geophysical techniques. Earthquake activity has focused attention on the central midcontinent, in the vicinity of He New Madrid seismic zone at the head of the .\lississippi Embayment, and has encouraged studies of contemporary tectonics as a means of predicting seismicity and the areal limits of He poten- tial seismic activity. This chapter reviews the major pub- lished tectonic hypotheses for the contemporary geody- namics of the midcontinent region. However? to set the firameworic for these hypotheses, the geological history is 73 summarized with emphasis on the structural develop- ment and related tectonic events. This summary is impor- tant because the tectonic events that have acted upon the midcontinent in He past are only interpretable based on the structural, sedimentary, and thermal events reflected in the geological history of He area and nearby plate mar- gins. If geological events involve orderly processes, then we can anticipate that the past in a general way is a clue to the present. Although noncyclic processes are impor- tant in early history and crystal conditions have changed to some degree, previous tectonism provides guidelines for subsequent dynamic processes (Allen, 1975~. It is im- portant to use this structural knowledge to decipher con- temporary tectonic processes. GEOLOGICAL HISTORY The geological history of the midcontinent region has been He subject of many discussions (e.g., Bristol and Buschbach, 1971), and the major tectonic events are shown schematically as a function of time in Figure 7.1. The early history of the area is poorly known because

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104 (w31..N . \J i ~-N .~- ~- N - -~Q FIGURE 9.7 Active and fossil hydrothermal systems in the western United States Dato ~~ we (1967). (a) Active wane c~- _~ ~ . - - ---- - ---- --~~~~~ -- -~ ~-~-- GORDON P. EATON' Eat _ _. F ~ generally are not compensated isostatically, suggesting, in tum, that faults serving as surfaces of adjustment do not pass through the lithosphere. All of these data tend to suggest that the depth limit of earthquakes may be Me depth limit of faulting. Stewart (1978) illustrated altema- tive interpretations of how such faults may tenninate downward. Figure 9.10 compares the aggregate ear~quake~epth distribution for the entire region with Basin ant! flange widths. The two histograms are generally similar in fonn, bow having intermediate values between O and 5 km peaking between 5 and 10 km, and having relative low levels of occurrence for values greater than 20 km. Of We earthquakes, 97 percent occur in the upper 15 km of the crust, which in We Great Basin Is We upper half of We crust. Of Basin and Range block widths, 88 percent are no wider than 20 km. Theoretical and experimental analyses of the depth and spacing of fractures torTned in extension (both tensile fractures and extensional shear fractures) suggest that the two values should be similar at least within an order of magnitude (LachenbIuck, 1961 Sowers, 19721. In a mechanical analysis of block gliding in which horsts and grabens formed, Voight (1973) denved an am proximate equation for the width of a block formed by extensional faulting of an initially continuous slab. It is or wretch are discussed in text W = 2T(45° - ¢/2), where W is the width of the block, T the thickness of the faulted slab; and ¢, the coefficient of Internal *iction for effective stresses. Application of this equation to the Basin and Range province requires We assumption of a surface or zone of translators sliding at the base of the fragmenting upper crust. It will be shown below that the Great Basin crust may have such a zone. In order to solve Voight's equation we must have a value for ¢. Byerlee's (1968) experimental study of the bnttle~uc~le transition in rocks indicated that Fiction is Independent of composition. Rocks under confining pres- sures of from O to 5.2 kilobars (~e depth equivalent of 0 to 16~5 km) revealed a remarkably systematic variation between increasing nonnal stress (a) and shear stress (T) for friction. The limiting slopes, eta = tan ¢, of Byerlee's finchon data curve are 36° and 46°. Substitution of these values in Voight's equation yields tl~e following wimps for fault blocks fanned in this manner: for shallow crystal slabs initially 10 km thick, widths of 8.1-10.2 km, for slabs 15 km thick, 12.1-15.3 km; and for slabs 20 km Wick 16.~20.4 km. These results are in good agreement win the observed Basin and Range block width~arthquake depth relation (Figure 9.10), hence Voight's model of ex- tensional sliding may be judged to have potential rele- vance to an understanding of the mechanics of Basin- Range faulting.

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Characteristics of the Crust of the Basin and Range Province Thompson (1959; 1966) was the first to suggest that extension in the deeper bas~n-range crust takes place via plastic stretching or injection of dikes. Hamilton and Myers (1966), Stewart (1971), and Proffett (1977) all ac- cepted the first of these concepts, viewing Basin and Range structure as fragmentation of ~ shallow crustal slab riding on a plastically extending substratum. Lateral dila- tion of We lithosphere by magma from below is implicit in the thermomechan~1 model of Lachenbruch and Sass (1978~. \\right and Trowel (19~.3) c~11ed upon both mechanisms (plastic stretching and intrusion) to extend Me deeper crust beneath Me fault-fragmenting surface slab of Me western Great Basin. NORMAL FAULTING AND BASAL SLIDING Laboratory-scale models of extensional faulting Mat use sand or dry mortar as Me deforming media (Hubbert, FIGURE 9.8 Seismicity of the west- em United States (from Smith, 1978). Lower threshold magnitudes were used in plotting California data. Heavy line, based on seismic data, marks in- board limit of highest earthquake event frequency and areal density in California, high cumulative seismic- strain energy release, and major strike- slip faulting related to dextral shear of the western plate boundary in Ho- locene time. Major faults within 150 km east of this line show oblique slip, with active stnke-slip components, but farther east, the dominant mechanism is simple extension. (Reprinted from Geological Society of America Memoir 152, with permission.) 105 1951; Stewart, 19~1) yield structures similar to single, simple grabens and distributed arrays of alternating grabens and horsts. Stewart's model, which was designed specifically to resemble Great Basin structure, had a sig- nificant feature—~ constructed surface of translatory slid- ing at its base. In order to predestine the spacing and plan of individual horsts and grabens, Stewart placed seg- mented sheets of paper beneath dry mortar. These sheets constituted a basal dislocation between the locally frag- menting mortar above and a sheet of uniformly extending rubber (the model's analog of a plastic substrate) beneath. Translatory sliding took place between the paper and the rubber sheet. In Hubbert's (1951) model, horizontal slid- ing took place at the base of the sand section, and it is not difficult to imagine striations developing on the floor of the deformation lynx parallel to the direction 'f e.xtensi<,n af;rer repeated runs. In a real earth, such a dislocation could take one of two forms: (1) a simple surface of sliding or (2) a Win zone of . . . . . . .. O _ .. , ~ . .. -I. . . · . . · - - . . _ . . . .:' ~ . . . . - . lo. . · . . ._ . C'~t' - -,\- ~= .~:_ a. · t~\ · . .. J A. . ..CC id- _ ,dz., · _ ~ · .: .- . · ·: · . . en do .. ' - - ~ EF~C£NTER MAP Of WEST thy UNITED STARS . ; . . . 45 ; . . ·. Cat . . . . .: By- 4~ . .... .

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106 ~ ~ _ DI '°~ W--~ ~ , ~ i- ~ I,lo. ~ /20 E ~.° 0 , ~ 0~~ °r 1° '50 , 20 1 J 0 so on 10~ 'art 20^ Q so FIGURE 9.9 Histograms of ear~qualce focal depths (in kilometers) for 14 areas within the region of active extension in the western United States. Data Mom Gumper and Scholz (1971); Jaksha et al. (1977); Ryall and Savage (1969); Shuleski et al. (197 7); and Smith (19783. listributed shear or decollement. A subhonzontal striated surface, a thin zone ofmylonite, or a thick layerofdynamo- thennally metamorphosed rock might Bus be anticipated, depending on kept, effective pressure, temperature, composition, and shear stress. The Turnagain Heights translatory slide in Alaska may be cited as an example of such clefonnation. It was de- scnbed and illustrated by Hansen (1965) and mechani- cally analyzed by Voight (19731. It mimics He Basin and Range province in structure. The ongina1 Sickness of He Augmenting slab was only 20 m, hence Here was little tendency for normal faults to flatten appreciably wig depth. Extension and sliding of He originally intact sur- face slab was possible because of an absence of lateral support on one side and a perceptible (2.2°) slope of the basal surface. Genetically, the structure was a gravity slide. It resulted from a loss in strength of material in He vicinity of He sliding surface as a result of excitation by a major earthquake. Owing to the backward retreat of its headwall, as more and more of the fragmenting surface slab slid away laterally, Voight termed He feature a "ret- rogressive block-glide." I do not suggest Hat the driving force of Basin-Range faulting is the same, only that kine- matics and resultant structures are grossly similar and, for this reason, instructive. Evidence suggests that the Great Basin may be growing laterally, i.e., that it may be consuming neighboring re- gions on both west and east (Smith et al., 1976; Eaton et al., 1978~. Both margins show transitional regions Hat GORDON P. EATON have geophysical anomalies characteristic of the Great Basin extending tens of kilometers into the neighboring provinces (see Ryall and Stuart, 1963; Shucy et al., 1973; Smith and Sbar, 1974; Keller et al., 1975; and Eaton et al., 1978~. The eastern margin of the Basin and Range prov- ince aIso has geological charactenshcs suggestive of tran- sition (Best and Hamblin, 1978; Howard et al., 1978; Luedke and Smith, 1978~. This state may relate, at least superficially, to the retrogressive aspect of He Turnagain Heights slicle. If so, the Simian and Wasatch fronts (opposed headwalls) may be retreating from each over as a result of plastic stretching at depth, Bus effecting a grown of He province at He expense of adjoining re- "ions. This could account for the obse~vecl outward re- striction of magrnatism with time. Uplift in the regions of these headwalls is probably a thermal phenomenon that is essentially contemporaneous wig the extensional fault- ing itself, just as it is in He oceans. Although heat flow in the Sier a Nevada is anomalously low, it must, in part, reflect the appreciable thermal time constant of He crust, for He mass deficiency characteristic of the Great Basin continues beneath the Sierra Nevadla (Eaton et al., 1978~. The Great Basin has a well-developed bilateral sym- met~y in certain aspects of its geology, but far more ob- viously in its geophysical fields (Proffett, 1977; Eaton et al., 1978~. In Proffett's model the western half of He shal- low, fragmenting, causal slab translates eastward relative to the extending substrate beneath (~e middle and lower crust), and Hat of the eastem half, westward. According to Voight (1973) retrogression cannot take place if die glicle blocks are need (as opposed to internally deformable) unless fluid pressures within fractures are sufficiently him to perform We fi'nchon of plastic wedges. Basin ~ Range blocks are sufficiently f~ctu red at We space ~ suggest internal deformation. We signifi- cant defonnationa] model thus appears to be lateral spreading win deformable block gliding. GEOPHYSICALLY ANOMALOUS LAYERS IN THE SHALLOW CRUST The possibility of a surface of sliding or a zone of ductile flow (mylonite or over metamorphic rocks) beneath He fault-fragmented surface slab of the Basin and Range province raises He question of Heir detectability by geo- physical means. We examine this issue only briefly, but in the last section of the chapter offer a tentative crustal model, based primarily on geology, heat flow, and ear~- quake dam, to which the other geophysical observations are fit by }~ypod~esis. The hypothesis needs specific test- ing. In He past decade and a half an increasing number of reports of a seismic low-velocity layer in He shallow crust have been published. One example is In He eastern Great Basin, near its boundary win the Colorado Plateaus (Mueller and ~ndisman, 1971; Landisman et al., 1971; Braile et al., 1974; Keller et al., 1975; Smith et al., 1975;

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Characteristics of the Cmst of the Basin and Range Province and Braile, 19771. Because such a feature is also asso- ciated with the Rhine graben (Landisman et al., 1971) it is tempting to conclude that it is a feature characteristic of extensional regimes. Shurbet and Cebull (1971) suggested that the crustal low-velocity layer in the Great Basin is a zone of de- creased rigidity that provides a means of absorbing the displacements of Basin-Range faults. According to them, the top of this zone is at levels of 5 km or so, and the base, at ~9 Icm. Braile et al. ( 1974) suggested that surface exten- sion by normal faulting at the surface is absorbed in a soft, plastically extending region of lowered seismic velocity. Braile (1977) later published the following additional information: (1) the layer has a compressional wave veloc- ity perhaps as low as 5.5 km sect (compared with veloci- ties above and below of 6.0 and 6.5 km see-i, respectively); and (2) it has a heightened Poisson's ratio; (3) an anoma- lously low Q (quality factor) for the transmission of com- pressional seismic waves; and (4) a top at 9.5 km and base at 15 km. Although the depth and thickness of this layer are different in the Shurbet and Cebul1 model, the values were derived in very different ways. The figures of Braile (1977) are preferred. Both sets of authors agree on the existence and gross mechanical properties of the layer, and on its role in accommodating extension at the surface. Neat as this picture is, it is marred by the fact that crustal low-velocity layers are not peculiar to extensional re- gimes. Two collections of papers on the physical properties and conditions of the continental crust (Heacock, 1971, 1977) reveal that the phenomenon is widespread, found in areas of young extension as well as in stable Precambrian shield areas (Berry and Mair, 1977~. Such features have been observed in the crust of all tl~e continents except Antarctica. Mueller (1977) has incorporated it as a key element in a generalized model of the continental crust. According to him, it is found fairly consistently at depths of ~15 km. Some investigators, however, place it as deep as 20 lam, e.g., Landisman and Chaipayungpun (1977~. Its origin has been ascribed to high temperatures (Smith et al., 1975), to Me presence of a zone of granitic intrusions (Mueller, 1977~, to high pore-fluid pressures (Berry and Mair, 1977), or to some combination of any or all of these factors (Mueller, 1977~. Laboratory experiments (Nur and Simmons, 1969; Todd and Simmons, 1972; and Brace, 1972b) demonstrate that as pore pressures rise toward li~ostatic values (thereby reducing effective pressure) seismic velocities fall toward those observed at exceedingly shallow levels in the crust. If a consensus as to the origin of the crystal seismic low- velocity layer is emerging, it is high pore pressure. In the Great Basin, high regional heat flow, which implies high crusty temperatures, probably plays an important sup- portive role. Pore water at high pressures in a closed system is capa- ble of lowering seismic velocity, Q values, and rock strength, conditions that appear to occur in the Great Basin crust. As Berry and Mair (1977) point out, however, 0 — s — 10 — 15— ye — 20 — 25 — 30 — 35— (A) 40— 10— IS— At 20— of 2S— 30— 35— (b) a_ 107 0 10 fREQUEl~Y. IN PEACES 20 30 So so 0 10 fREQUEI=. ~ ~~ 20 30 40 50 · I l l l ~ ~ , ~ ~ , . . . FIGURE 9.10 Earthquake focal depths and widths of Basin and Range blocks. (a) Histogram offocal depths of 2,4~o earthquakes in the region of extension; (b) widths of individual basins and ranges in the Great Basin scaled from the map of King and Beik- man (1974). Both characteristic dimensions (depth and width) show modal values in the range ~10 km, with most values (>85 percent) in the range ~20 km.

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108 this explanation requires that rocks in the layer in ques- tion have finite porosities at depths of ~15 km, while those in the zone immediately above it must be free of permeability because the hydraulically pressured layer must have some sort of impermeable cap. This cap in the Great Basin may be a layer of rock extended by ductile flow, the one whose upper surface may mark the base of the region of brittle faulting. If so, the seismic low- velocity zone may not coincide with this layer but is per- haps beneath it. Depths to surfaces or zones of Tertiary sliding in the Great Basin have been estimated by Arm- strong (1972) to be at least 8 km on geological grounds, but Me Wick, ductile zones must be deeper still, because in many places the glide faults do not rest directly on ductile rocks. Because depths of somewhat more than 8 km are in reasonable agreement with Braile's (1977) seismic es- timate of 9.5 km to the top of We seismic low-velocity layer, I tentatively regard the ductile layer (if it tally exists) as a possible cap. Pore fluids could be trapped at high pressure in fractures in rock immediately beneath such a ductile layer. Internal displacements or structural adjustments within rocks at this level would take place by stable sliding (Brace, 1972a; 1972b), and earthquakes would not be common at such depths. As we have already seen, they are not. Pore water in closed-rock systems will also lower elec- trical resistivity, as will increasing temperature' which elevates the ionic mobility of such fluids. At very him temperatures the onset of partial melting could do much the same thing; the melting temperature of the rocks is lowered by the presence of water. For these reasons one might anticipate the presence of electrical conductors in the Great Basin crust, and, in fact, they are observed. Some investigators (e.g., Landisman and Cha~payung- pun, 1977; Lienert ant! Bennett, 1977) have equated the crustal low-velocity layer with a low resistivity layer in tectonically active or high heat-flow areas. Much remains to be done to substantiate this equivalence, and also to establish equivalence between a subhonzontal low- velocity layer in the crust and a porous zone capped by a stratum of ductile impermeable rock. The data in hand are permissive, but thus far hardly conclusive. The electrical data are reviewed briefly to provicle an idea of what is currently known. I am indebted to my fnend and col- league, I. N. TowIe, for providing the information sum- mary that follows. Schmucker's (1970) geomagnetic variometer investiga- tions in California indicated the presence of an electrical conductor in the western Cordillera at the eastern base of the Sierra Nevacla, at and near its boundary with the Great Basin. Stanley et al. ( 1976b) identified a shallow (2~7 'lan), highly conductive layer in the crust beneath the Carson sink in western Nevada by means of magnetotellunc soundings. Stanley et al. (1976a) also studied the elec- frical structure of the Long Valley geothermal system in the western Great Basin by means of direct current and electromagnetic techniques, concluding that hydro- ~ertnal activity is reflected in discrete conductive zones GORDON P. EATON in the crust, which are controlled, in turn, by regional faulting. Lienert and Bennett (1977) have identified ~ crystal conductor in the western Great Basin at a depth of 90 lam using controlled-source geomagnetic vanomet~y. Reitzel et al. (19703 and Porath and Cough (1971) ob- served generally reduced vertical geomagnetic field vari- ations in the eastern Great Basin Mat Hey interpreted as reflecting a shoaling of Me mantle. Ambiguities in Weir interpretation of crustal thickness will doubtless be re- solved as the evidence mounts both for a shallow, strongly conduchng, crustal layer in the Great Basin as a whole and for local shoaling of the asthenosphere. Studies by W. D. Stanley and colleagues at the U.S. Geological Survey (personal communication, 19~8) have revealed the presence of ~ conductive crustal laser near the boundary between the Great Basin and Snake River Plain on the norm. Depths to its top range from 2 to 10 km; its thickness may be as great as 10 km. Stanley et al. (1977) have also identified a conductor beneath the Snake River Plain region at depths of only 5 lam in the Yellowstone caldera, but deepening to 20 km on Me southwest. In the vicinity of the Raft River geothermal area, in the northeastern Great Basin, it is 7 km deep. This conductor may be related directly to the presence of magma, at least in the Yellowstone area, and, therefore, may or may not be directly related to the crustal low- velocity layer under discussion. On the basis of these limited data it appears that an electrically conductive layer is a common feature of He shallow Great Basin cmst. Depth estimates place its top between 2 and 20 kin, and in several areas, at less than 10 km. Possibly this conductor coincides with the crustal low-velocity layer, but too little is known about it to be certain. Coincidence might be anticipated simply be- cause some of the factors that lower seismic velocity also raise electrical conductivity (high temperature, high porosity, the presence of a pore fluid, or the presence of a silicate melt). An electrically conductive layer by itself does not require abnormally high pore pressures and, hence, does not require the presence of an impermeable cap to keep the system closed. The presence of conduc- tive minerals such as metallic sulfides or those having high ion-exchange capacity, like clays or zeolites, can also lower rock resistivities without affecting seismic velocity. The low-resistivity layer in the crust could just as well be above the low-velocity layer, reflecting some combination of high porosity, temperature, pore-fluid salinity, or hy- drothertnal alteration in the lower part of the shallow crust. CRUSTAL MODEL FOR THE BASIN AND RANGE PROVINCE: A SUMMARY AND INTERPRETATION The crust of the Great Basin section of the Basin and Range province (and its immediate environs) is higher in elevation, thinner, warmer, more highly fractured. and

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Characteristics of the Crust of the Basin and Range Province more well endowed with hot springs than that of sur- rounding regions, excluding the area immediately north of Me Snake River Plain. The fractures (mostly faults) extend a third of the way to halfway through Me crust. They are loci of abundant shallow earthquakes and vigor- ous hydrothermal circulation. Cnlstal extension takes place by faulting near the surface, but probably takes place by other modes at depth, most likely by dike intru- sion and stretching of Me lower crust and lithospheric mantle, and by ductile shear flow (distributed decolle- ment) in a relatively thin layer at some intermediate level. Repeated magmatic invasions of the crust have taken place during the past 100 million years. Some of these magmas broke the surface, but some have come to rest within the crust' giving up their heat Mere. Shallow mag- matic systems serve to drive hydro~ennal convection in the shallow crust, as does high regional heat flow from Me deeper crust. The phenomena have been long lived. At present, extension is taking place in an east-west or west-northwest direction; earlier, it was directed sou~- west or west-southwest (Eaton et al., 1978; Zoback and Thompson, 1978~. The Sonoran Desert was included in the initial episode of extension but is not included in the present one. Because the Sonoran Desert is deeply eroded, it exposes Me effects of crustal extension at deeper levels. A large part of the Sonoran Desert in south- western Arizona reveals evidence of subhonzontal, unidi- rectional plastic strain of middle Tertiary age (Davis, 1977; Davis et al., 1977; Davis, see Chapter 8; Rehrig and Reynolds, 1977) that initially developed before block faulting began but that may have served as the base of We faulted' shallow slab. The mylonitization and metamor- phism probably took place at lithostatic pressures of several kilobars. Armstrong ( 1972) reviewed evidence of transIatory dis- placements in the Basin and Range province and argued Mat some of the subhonzontal surfaces of sliding in the Great Basin are certainly Tertiary in age, Mat many of them may be Tertiary, and that they are more likely related to Basin and Range faulting Man to the Sevier orogenic event' the youngest episode of pre-extension thrusting. Most of these dislocations place younger strata over older. He noted that some of the structures are of relatively deep-seated origin (at least 8 km). It is possible that these subhorizontal zones of sliding first developed as thrust soles during crustal compression, later to evolve into extensional decollements. These observations and speculations lend themselves to the interpretation that prolonged thermal conditioning of die crust plus hori- zontal shearing simply may have continued earlier ini- tiated dynamotherrnal metamorphism of rocks Mat now serge as a boundary layer between parts of Me lithosphere extending by fundamentally different mechanisms. As young nominal faults developed near the surface in the regime of crustal extension, they became Iistric to (they came to sole on) older thrust zones. The mechanical model of Kehle (1970), in which ~ decollement is distributed through the middle (relatively 109 m're ductile) leaver Elf ~ cn~.st~1 fir lith~;pheric triad. my I'e applicable. Shearing in such a layer would lead to the mechanical generation of heat. Its magnitude w ould be controlled by the rate of shearing, which, judging from rates of extension measured at the surface, should be lower than that generated along the San Undress Fault (see Lachenbluch and Sass, 19 31. S'me part <'f ~e high heat loss in the Great Basin could be due to shearing, however. The greater part has been ascribed to penetra- tive convection of Me lithosphere by basaltic magma (Lachenbruch and Sass, 1978~. Part is ascribable to con- vective groundwater circulation in the shallow crust and locally, to young, hot, volcanic systems residing in the upper crust. If this model, based largely on geological, heat flow, and earthquake data, is generally correct, it has implications for the surface patterns of deformation, the regional distribution of geological resources, and some of the effects of earthquakes. The model is shown in sche- matic fore in Figure 9.11. To summarize its implications: 1. The location and extent of the Basin and Range prov- ince may have been largely predetermined by Me loca- tion and extent of early Tertiary and Mesozoic magmatism that preheated (thermally weakened) the crust and augmented a regime of compressiona1 thrusting in which subhonzontal dislocation surfaces fir ductile zones (dis- tributed decoIlements) first developed at middle to shal- low crustal levels. Such zones could be used later as basal dislocations for normal faulting and might also serve (at depth) as crustal membranes impermeable to the deeper circulation of groundwaters but allowing the upward passage of magma by intrusion. Normal faults at the sur- face probably are listric to Me deepest of these zones, inasmuch as the shallowest ones are exposed in Me uplifted fault blocks themselves. 2. Zones of translatory, ductile shear would constitute near-horizontal surfaces of mechanical decoupling or in- e~cient coupling within Me crust. As a result, deforrna- tions en cl kinematic motions in the lower crust would not always be clearly or faithfully reproduced at We surface. 3. The regional maintenance of long-continued high temperatures and high permeability assures the continua- tion of vigorous hydrothermal circulation and attendant epigenetic deposition of minerals through both compres- sional and extensional regimes. They may' on the other hand, be responsible for what appears to be a regional scarcity of oil and gas in the Basin and Range province. Where unfavorably situated, such fluids could be driven to the surface, where they could escape, except for local conditions of entrapment. For the same reason, considera- tion of Great Basin sites for the isolation and storage of radioactive wastes cames with it Me requirement of a critical evaluation of the local hydrologic and seismo- tectonic regime. 4. Seismic energy traversing the crust of the extended region is probably absorbed to a somewhat greater degree Man it is in the relatively less intensely fractured, shallow crust of the central and eastern United States, hence the

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110 GORDON P. EATON 1~.'~ If. i,,, At' W: I L-1 ,~L 1 .-...,,__.' L-3 L-4 FIGURE 9.11 Interpretive model of possible crustal structure dynamothermally metamorphosed Miocene and older rocks of a of the Great Basin (simplified, schematic, and not to scale); based wide varied of original compositions (medium stippling). At one on surface geology, heat flow, and earthquake distribution. (See extreme, the layer is a vanishingly thin stratum of mylonite, 1-10 Stewart, 1978, for alternative interpretations of the near-surface mm thick; at the other, a layer of granitic augen gneiss, schist, or structure.) The crust is composed of three layers (L-1, L-2, L-3) amphibolite, 1-3 km thick. This layer is locally or regionally having different lithologies and physical properties. Each fails or lineated and extends by laminar plastic flow. It is generally im- yields in extension by a different physical mode. Erosion in the permeable to groundwater circulation except where later Sonoran Desert region generally has cut down to the level of L-2. uplifted and fractured in the brittle regime, but at depth it may be LO is lithospheric mantle. Characteristics of these layers are as cut by dikes of Tertiary igneous rocks (solid block). It developed follows: L-1, Fault-fragmented, surface layer, ~15 km thick, first as a regional thrust sole (heavy dashed line) during earlier composed of rocks of a great variety or origins, compositions, and crustal compression. L-3, lower crustal layer, 10-20 km thick. ages, all exposed at the surface somewhere in the region; diagram composed near its top of igneous and metamorphic basement shows Cenozoic continental sedimentary and volcanic rocks at rocks like those of layer ~l but grading downward into increas- the surface (patterns of dense stippling and solid black, ing proportions of old granites, migmatites, gneisses, amphi- respectively)overlyingoldersedimentaryrocks(openstippling) bolites, and felsic to mafic granulites, in approximately that and granitic and metamorphic basement rocks (plain white). Al- order. This layer extends by a combination of diking (by basalts though the diagram does not show it, stratified rocks extend well from the asthenosphere, solid black) and solid-state convection into L-2 and probably into L-3 (as in Arizona). All of these rocks (stretching and underplating). It is rigid at relatively high and are highly fractured, as indicated by the plexus of fine, irregular intermediate strain rates. These modes of penetrative convection lines. The layer fails in semibriKle fashion by normal faulting, are responsible for the mass transport of heat from the deepe fault-block rotation, pervasive fracturing, and slumping. The mantle, causing the anomalously high heat flow observed in the deformation creates high fracture porosity and permeability, al- province. The seismic low-velocity zone may coincide with the lowing convechve circulation of groundwater (curved arrows) uppermost part of this layer as a result of anomalously high pore driven both by the high heat flow from the deeper crust and by pressure in a system capped by the impermeable layer, L-2, and local, young intrusions (black, dike-like bodies). The upper part high temperature. Let, lithospheric mantle. 25-35 km thick, of the crustal low-resistivity layer may coincide with the lower composed of ultramaf~c rock devoid of finely disseminated melt. part of this layer. The base of the layer generally marks the man- This layer, like the lower crustal layer above it, extends by diking imum depth of earthquakes. L-2 ductile intermediate layer, 0-3 (perhaps via rising, bleb-like bodies) and solid-state convection. km thick, composed of pervasively sheared, mylonitized, and/or It is immediately underlain by asthenospheric mantle.

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Characteristics of the Crest of the Basin arid Range Province geographical extent of isoseismal boundaries for an ear~- quake of given magnitude is generally less in the West than it is in the M~clwest ant! East. ACKNOWLEDGM ENTS Preparation of this review was aided substantially by dis- cussions win, and/or data con~ibudons from, M. D. Cnt- tenden, fir., G. A. Davis, Spiel Friedman, W. Hamilton, A. H. Lachenbruch, P. W. Lipman, E. H. McKee, S. S. Oriel, A. R. Sanford, B. B. Smith, T. A. Steven, and J. N. Towle. The manuscript benefited from constructive re- views by K. A. Howard and H. M. Iyer. None of the above should be regarded as subscribing to the interpretive crustal model presented here, however, nor to its stated implications. I am indebted to members of the technical staffofthe Hawaiian Volcano Observatory, Ron Hanatani, Rick Hazlett, Jenny Nakata, and Maunce Sako, as well as summer student aide Kevin Cuff for graphical comp~la- tion of many of the regional geological and geophysical data presented here. REFERENCES Anderson, R. E. (1971). Thin skin distension in Tertiary rocks of southeastern Nevada, Geol. Soc. Am. Bull. 82, 43~8. Armstrong, R. L. (1972). Low-angle (denudation) faults, hinter- land of the Sevier erogenic belt, eastern Nevada and western Utah, Geol. Soc. Am. Bull. 83, 172~1754. Atwater, T. ( 1970). Implications of plate tectonics for the Cenozic tectonic evolution of western Now America, Ceol. Soc. Am. B?~11. 81, 3513~536. Ba~angi, M., and J. Dorrnan (1969). World seismicity maps compiled from ESSA, Coast and Geodetic Survey, epicenter data, 1961-1967, Bull. Seismol. Soc. Am. 59, 369 380. Berry, M. J., and J. A. Mair ( 1977). The nature of the earth's crust in Canada, in The Earth's Crust, J. G. Heacock, ea., Am. Geo- phys. Union Geophys. .\lonogr. 520, pp. 319~348. Best, M. G., and W. IC. Hamblin (1978). Origin of the northern Basin and Range province: implications from the geology of its eastern boundary, in Cenozoic Tectonics and Regional Geo- physics ofthe Western Cordillera, R. B. Smith and G. R. Eaton, eds., Ceol. Soc. Am. Mem. 152, 313~340. Brace, W. F. (1972a). Laboratory studies of stick-slip and their application to earthquakes, Tectonophysics 14, 18~200. Brace, W. F. (1972b). Pore pressure in geophysics, in Flow and Fracture of Rocks, H. C. Heard, I. Y. Borg, N. L. Carter, and C. B. Raleigh, eds., Am. Geophys. Union, pp. 26~273. Brace, W. F., and l D. Byerlee (1968). Stick slip, stable sliding, and earthquakes- effect and rock type, pressure, strain rate, and stiffness, J. Geophys. Res. 73, 6031~37. Braile, L. W. (1977). 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