All Men Dinger 1990

TECTONICS, VOL. 9, NO. 4, PAGES 789-809, AUGUST 1990

FORELAND SHORTENING AND CRUSTAL BALANCING IN THE ANDES AT 30S LATITUDE

R. W. Allmendinger,1 D. Figueroa, 2 D. Snyder,1 J. Beer, 1 C. Mpodozis, 3 and B. L. Isacks 1

Abstract. Excellent surface exposures, known Benioff zone geometry, a dynamic morphology, and the availability of industry seismic reflection data all make the Andes at 30S an excellent transect for investigating crustal-scale balanced sections. 150-170 km of horizontal shortening has occurred in three major belts located between the trench and the foreland. The thin-skinned, east-verging Precordillera of western Argentina accounts for 60-75% of the total shortening and formed mostly since major volcanism ceased at -10 Ma. Industry seismic reflection data show that the d(collement of the Precordillera belt is located anomalously deep at -15 km. The belt is dominated by fault propagation folds and contains several prominent out-of-sequence thrust faults. Seismic stratigraphic analysis shows that Miocene strata in the Iglesia

1Department of Geological Sciences, Cornell Univer- sity, Ithaca, New York; Snyder now at BIRPS, Univer- sity of Cambridge, Cambridge, U.K.; Beer now at Pecten International Company, Houston, Texas.

2yacimientos Petroliferos Fiscales, Buenos Aires, Argentina.

3Servicio Nacional de Geologia y Mineria, Santiago, Chile.

Copyright 1990 by the American Geophysical Union.

Paper number 89TC03487. 0278-7407/90/89TC-03487510.00

Valley, located between the Precordillera and the crest of the Andes, accumulated in a piggy-back basin. Onlap relations on the western side indicate that the High Cordillera was uplifted as a major fault bend fold over a buried ramp. Thrusting in the two western belts, both in the High Cordillera of Chile, formed during the waning stages of arc volcanism, 11-16 Ma. and account for 2540% of the shortening. The observed shortening is probably greater than can be accounted for with reasonable crustal thicknesses, indicating the possibility of continental truncation or erosion along the plate margin or an anomalously thick root held down by the nearly flat subducted Nazca Plate. Our preferred crustal geometry puts the ramp between upper and lower crustal deformation west of the high topography, requiring crustal scale tectonic wedging to thicken the crust beneath the crest of the Andes. This non-unique model provides a simple explanation of the first order morphology of the Andes at this latitude.

INTRODUCTION

Balanced cross-sections have proven very powerful for understanding that part of orogenic belts closest to the undeformed craton: foreland thrust belts. More difficult, however, is the documentation of the kinematics of entire orogenic belts via the same approach. Commonly, there are two limiting factors: (1) magmatic addition to the crust is ;. "ild card" that can seldom be quantified, and (2) the "boundary conditions" pertaining to that end of the cross section which is not pinned to the craton are seldom known because of cover, complex multiphase structures, strike-slip faulting and/or ductile deformation.

In the Andes at 30S latitude (Figure 1), these constraints are less serious than elsewhere. This segment

790 Allmendinger et al.: Shortening in the Andes at 30S

12

16

20

24

28

32

i i

0 km ]000

78 74 70 66 62

Fig. 1. Regional morpho-tectonic map of the Central Andes [from Isacks, 1988], showing the location of the study area in (depicted in Figure 2). Contours show depth to the Benioff Zone, in km. Gray shaded area in South America is that part of the Andes above 3 km elevation. Horizontal ruled pattern shows the extent of the Sierras Pampeanas of western Argentina.

of the Andes has been largely amagmatic for the last ~9 Ma [Kay et al., 1988], although younger, volumetri- cally insignificant volcanism is present locally. The deformation addressed here began at ~16 Ma and thus overlaps in time with that magrnatism; the majority of shortening in this segment occurred during the waning phases or post-dates the magrnatism [Jordan et al., 1988; Beer and Jordan, 1989; Damanti and Jordan, 1989]. Further advantages of the Andes at 30S include excellent surface exposures due to the arid climate, the availability of industry seismic reflection data (presented below), and well-known crustal and Benioff zone seismicity. The latter is particularly important because the subducted plate provides the ultimate boundary condition for any attempt at crustal scale balancing. Finally, the orogen is relatively narrow in this segment with an antithetic thrust belt (the Precordillera of western Argentina) located unusually close to the plate boundary (

Allmendinger et al.: Shortening in the Andes at 30S 791

Pacific

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:Argentina + + .... :.:.:.:.:.:.:.:.:.:.:.:.:.:.:.:.:.:.:.:.;.:.;.:.:.. Bermejo

Basin

Fig. 2. Generalized tectonic map of the Andes at 30S latitude, showing the major belts and structures described in the text. The diagonal line shows the location of the cross-sections in Figures 3, 15, 16, and 17. Light gray boxes show the location of seismic reflection data available for this study (Figures 5, 6, and 10). White, Quaternary basins; dashes, Sierras Pampeanas; stipple, outcrop of upper Cenozoic strata; tilted "bricks" (limestone), Precordillera structural province; small pluses, Permo-Carboniferous strata and intrusions of the Argentine Frontal Cordillera; "v's", Miocene volcanics along the international border; large pluses, Permo-Carboniferous crystalline core of the Chilean High Cordillera; gray shade, Mesozoic and lower Tertiary rocks of the Chilean Coast Ranges.

seismicity to about 40 km [Kadinsky-Cade et al., 1985; Smalley et al., 1988]. Precordillera

The Precordillera is a thin-skinned thrust belt which forms the "foothills" of the physiographic Andes between 29 and 33S (Figure 2). Typically, four to six major west-dipping thrust faults comprise the belt depending on latitude. Most of the thrusts apparently have their dEcollement level in a Cambro- Ordovician limestone sequence [Baldis and Chebli, 1969; Ortiz and Zambrano, 1981]. The westernmost thrusts expose Ordovician to Devonian flysch, a slope facies at least partly equivalent in age to the limestone [Baldis et al., 1982]. Although Paleozoic thrusts and fold axes in the west part of the belt dip eastward, Andean-aged thrusts consistently dip to the west.

The age of Precordillera thrusting has recently been constrained by extensive chronological and geological studies of foreland basin strata east of and within the belt [Jordan et al., 1988; Beer and Jordan, 1989]. At 30S, these data indicate that thrusting began in the limestone part of the belt at -9 Ma and in the western

part of the belt at 14-16 Ma. Approximately 100 km to the north, thrusting probably began at -16 Ma [Reynolds et al., 1987b; Damanti, 1989] and new geochronology indicates similar antiquity at 30S [Beer et al., 1990].

The western margin of the Precordillera is formed by the Iglesia Valley (Figure 2), which is underlain by a 4-km-thick basin of upper Tertiary strata (~ 15 Ma and younger; Beer et al., [1990]). As shown below, the Precordillera accounts for -65-70 % of the total shortening in this segment of the Andes. The structure of the Precordillera and Iglesia basin together comprise a major focus of this paper.

High Andes

West of the Iglesia basin, the Andes rise to elevations of -6 km along the international border between Chile and Argentina. Farther south, the High Cordillera can be divided into the Frontal Cordillera to the east and the Principal Cordillera to the west. The thick sequence of Mesozoic and Tertiary volcanic and sedimentary strata that characterizes and is restricted to the latter, however, crops out discontinu-


ously northward and there is no clear physical boundary between the two belts at the latitude of the transect. On the eastern slope (Figure 2), the High Cordillera exposes gently-dipping, openly-folded Carboniferous-Lower Permian marine strata, Late Paleozoic-Triassic granitoid batholiths, and Permo- Triassic ignimbrites of the Choiyoi Group [Caminos, 1979; Ortiz and Zambrano, 1981; Llambias et al., 1987]. No major belts of Andean shortening are present between the Iglesia Valley and the international border. An imbricate belt of thrust faults has been mapped immediately north of the northern termina- tion of the Iglesia Valley [Marin and Nullo, 1988]. The fact that none of these structures are observed to cut the Neogene deposits of the Iglesia basin suggests either that they are older than mid-Miocene in age or that a major, presently undocumented east-west structure at the northern end of the valley has produced north-south segmentation of the Andes at these latitudes.

A belt of east- and west-dipping thrust faults, known as the E1 Indio belt, straddles the international border [Maksaev et al., 1984; Ramos et al., 1990]; at 30S it is almost entirely within Chile but to the north and south it passes into Argentina (Figure 2). These structures deform Mesozoic and Tertiary volcanic units and were active until --16 Ma (a major unconformity exists between the Dofia Ana Formation, 29 - 18 Ma, and the Cerro de las T6rtolas Formation, 16.6 - 11 Ma). In Argentina northwest of the Iglesia Valley, the Cerro de las T6rtolas Formation locally is also tilted up to 35 and is unconformably overlain by the undeformed 6 Ma Vacas Heladas ignimbrite [Ramos et al., 1990]; the regional importance of this deformation is unknown but is interpreted here to be small. The western border of the E1 Indio belt is the west-dipping Barios del Toro fault which uplifts the eastern side of the Paleozoic plutonic nucleus of the Chilean High Cordillera. The eastern border is formed by an east- dipping thrust system in the Valle del Cura of Argentina [Marin and Nullo, 1988].

The plutonic nucleus of the Chilean High Cordil- lera is a 40-60 km wide belt of Late Paleozoic-Triassic

composite batholiths which extends for more than 350 km along strike between 31 and 28S [Cornejo et al., 1984; Nasi et al., 1985]. A substantial amount of shortening, mainly pre-Late Miocene in age, is recorded in this region particularly north of 30S. Several "gh-angle" reverse faults can be traced for more than 80 km along strike and basement cored anticlines attest to distributed deformation within the major fault blocks [Moscoso and Mpodozis, 1988]. Many fewer faults have been mapped at 30 but this may be due solely to the deeper erosion and lack of supracrustal strata which farther north facilitate the identifi- cation of structures. Alternatively, there may be north-south changes in the distribution of shortening, as suggested earlier.

The western side of this plutonic nucleus is formed by the Vicufia thrust fault (Figure 2). This structure dips 25-45 east and can be traced for a strike length of

230 km [Dedi6s, 1967; Moscoso and Mpodozis, 1988]. Throughout this distance, the thrust juxtaposes the Paleozoic plutonic belt against Cretaceous and lower Tertiary volcaniclastic strata of the Coastal Cordil- lera province. In fact, this fault may be localized along the the old eastern margin of the Cretaceous volcano-tectonic rift of the Chilean coastal region. In the hanging wall, a 15-km-wide belt of Mesozoic volcaniclastic strata (Figure 2) are deformed into an open syncline (the Guanto syncline, [Mpodozis and Cornejo, 1988]) which parallels the trace of the thrust. The Vicufia fault clearly correlates with a distinct, first-order break in the regional topography but it is unclear whether this step results from recent movement or simply different erodability of lithologic units on either side of the fault. The age of the last and largest movement of the thrust fault is constrained stratigra- phically only to be younger than early Tertiary. Cross-cutting relations between early Tertiary intrusions and cleaved limestones in the Hurtado region suggest that the fault-syncline couple may actually have begun to form in Late Cretaceous time.

Coastal Region

To the west of the Vicufia fault, the topography is much more subdued. More than 5 km of Cretaceous volcaniclastic rocks, covered locally by lower Tertiary volcanics and intruded in the west by a long belt of Cretaceous batholiths, occur in this region [Thomas, 1967; Moscoso et al., 1982]. The Cretaceous section represents the infill of a volcano-tectonic extensional depression which may have formed immediately prior to the modern compressional history of the Andes [Mpodozis and Ramos, 1990]. There is little evidence of significant shortening in the Coastal Region; Creta- ceous strata generally dip gently eastward with a single broad anticline at the center of the belt [Aguire and Egert, 1965]. Major strike-slip faults have not been identified in this transect of the Chilean Cordillera although they are known both to the north and south.

FIRST ORDER BOUNDARY CONDITIONS

Topography

In an active mountain belt, topography is a dyna- mically evolving feature which directly reflects first- order tectonic processes [e.g., Isacks, 1988]. On casual inspection, the Andes at 30S appear to rise smoothly from the trench to the crest (Figure 3). East of the crest, the gradual reduction in topography is inter- rupted by the Iglesia Basin-Precordillera couple. Sig- nificantly more detail can be gleaned from a hypsometric curve along the same transect (Figure 4a) because 1:1 or even 5:1 topographic profiles of an entire mountain belt, reduced to page size (e.g., Figures 3, 4b) show virtually none of the detail of the original data set. The differential hypsometric curve, in particular, identifies an important surface at --4 km elevation, which is also clearly visible in a profile of maximum

Allmendinger et al.- Shortening in the Andes at 30S 793

Coastal Region

Average raphyi'

5o km

High Cordillera Iglesia Basin Precordillera

11ement level ........................

Sierras Pampeanas

Fig. 3. Outline sketch of lithospheric-scale boundary conditions as described in the text. Stippled area on the east side is the Bermejo foreland basin. Benioff Zone is from Cahill and Isacks [1985] and Isacks [1988].

elevations (Figure 4b). Thus the crest of the Andes at this latitude constitute a narrow plateau, approximately -80 km wide. Peaks rise above this surface to elevations of 6.0 km and the surface is also dissected by canyons, particularly on the west side (Figure 4b).

The timing of uplift of the Andean edifice is a con- tinuing controversy, and no direct evidence bears on the age of morphological development at 30S. On a more regional scale, the present Andean morphology is thought to be largely the result of mid-Miocene and younger uplift [Pascual and Odreman Rivas, 1973]. Since the mid-Miocene, the High Cordillera appears to have been regionally uplifted without significant internal shortening, whereas the Precordillera has experienced active shortening. This result matches well with the unroofing history as recorded in the upper Cenozoic basins at 30S [Beer et al., 1990; Damanti, 1989] and with the timing of crustal thickening inferred from geochemical studies [Kay et al., 1988]. We use the narrow plateau surface in our interpretation of the structure beneath the High Cordil- lera. The model presented below suggests, but does not prove, beginning of uplift of the High Cordillera at 12-16 Ma.

Benioff Zone

Just as the form of the free surface provides an upper boundary condition, the subducted Nazca plate must ultimately provide both the lower and the western limits to a crustal-scale balanced section (Figure 3). Considerable data from both local and world-wide networks constrain the geometry of the subducted plate [Barazangi and Isacks, 1976; Bevis and Isacks, 1984; Cahill and Isacks, 1985; Smalley and Isacks, 1987; Isacks, 1988]. These data clearly show that the Benioff zone has a normal, 30 east dip in the

interplate zone but at depths below 75-100 km it has an unusually low dip (-5).

The low dip has two important, related effects: a full thickness of lithosphere is missing beneath Chile and western Argentina, and more importantly, there is little or no asthenospheric wedge between the two plates. The shallowing of the plate between 10-15 Ma [Kay et al., 1988] cut off significant magmatism at about that time and thus removed it as a major factor in subsequent crustal volume change. The lack of a significant asthenospheric wedge means that the present topography of the region probably does not have a strong thermal component [Isacks, 1988]. Crustal Seismicity

Although crustal seismicity does not provide a physical boundary condition, it does shed some light on both crustal rheology and depth distribution of brittle shortening. Crustal earthquakes in the westernmost Sierras Pampeanas occur to 40 km depth, only 70 km above the Benioff zone and


A. [ [ [ ] [ I [ [ [ [ [ I 200 100 I 100 200 300

- WEST Area / km 6'0 i!.?....... Area / km EAST -

Cummulative Area (kin) I Cummulafive Area (kin2) / continental divide /

B. max topo plateau surface ave topo\ ' Iglesia

Fig. 4. (a) Hypsometric curves of the Andes at 30-32S. Irregular curve with gray shaded area (top axis) is the incremental (differential) curve which shows the amount of surface area at any particular elevation. Note the maximum in area at 4 km elevation. The smooth gray curve (bottom axis) is the cumulative curve which shows the total area above a particular elevation. It is the integral of the incremental curve. Compare the form of these curves to those shown in Figure 2 of Isacks [1988] for the Altiplano segment of the central Andes. (b) Topographic cross-section of the Andes along the line of section in Figure 2, with elevations 20 km on either side of the line of section projected onto the line. The "max", "ave", and "min topo" have 5x vertical exaggeration. Section is generalized due to both projection of a swath of data and the very reduced scale of the display on a base showing the curvature of the Earth. The plateau surface is best seen on the maximum topography curve, the incision of rivers on the minimum curve. The difference between the two shows the degree of erosion. Note the much more extensive erosion on the west (Chilean) side of the Andes. This is also reflected in the hypsometric curves.

Inferred Crustal Thickness

The most poorly constrained part of this analysis is the lack of any direct measure of crustal thickness. No crustal scale refraction surveys exist anywhere in western Argentina, and extremely limited deep reflection data do not image the base of the crust. Even gravity

data are not available. The only gravity surveys that cross the orogen are located >250 km to the south [Introcaso, 1980] and >500 km to the north. Thus, we can make only simple isostatic models which are consistent with the topography (Figure 3). Unfortu- nately, even the degree of local versus flexural compensation is unknown. The extent to which our crustal

Allmendinger et al.' Shortening in the Andes at 30S 795

thickness models are reasonable can be gauged by how closely observed surface shortening matches that predicted using crustal scale area balancing. This sort of reasoning cannot be pushed too far, however, without encountering its inherent circularity.

Depth to the Dcollement

One of the most fundamental controls on the bulk shortening in a thin-skinned deformation system such as the Precordillera is the depth to the d(collement. Industry seismic reflection data and the surface geology in western Argentina provide very good control on this vital datum. It has long been noted that the Cambro-Ordovician limestone or flysch sequence occurs at the base of each major thrust plate in the Precor- dillera; older rocks are never exposed [Baldis and Chebli, 1969; Ortiz and Zambrano, 1981]. This strongly suggests that the regional d(collement lies within that sequence and above basement. The pre-Cenozoic stratigraphic thickness above that level varies considerably due to a complex, pre-Andean structural history. In general it is about 4 km on the east side of the Precordillera and 6 km in the middle [Furque, 1979;

Fielding and Jordan, 1988] (also see below). Farther west, the original thickness of the Cambro-Ordovician flysch in unknown because of the intense, pre-Andean folding of that part of the sequence.

Recent seismic reflection profiles collected for Yaci- mientos Petolfferos Fiscales (YPF) cross the front of the easternmost west-dipping thrust (referred to below as the Niquivil thrust). Those data (Figure 5) show dipping reflectors within the thrust plate overlying prominent flat lying reflectors at 4.5 and 5.1 s (two- way time). The flat reflectors cannot be multiples and the line geometry with respect to all known three dimensional features argues strongly against a side- swipe interpretation. Given what is known of the velocity structure from the stacking velocities, we calculate that the two horizontal reflectors lie at

and --15 km depth. Given the vertical separation of these reflectors and their continuity, we suggest that they correspond to the top of the Cambro-Ordovician limestone sequence and the top of basement, respectively. These depths are also consistent with the km thickness of Miocene and younger synorogenic strata in the foreland basin directly east of the easternmost thrust (Figure 6). Thus, the basal d(collement at the

W

out-of-sequence thrust

E

Niquivil thrus

Fig. 5. YPF seismic reflection data from the Niquivil thrust (arrows), showing the depth to d(collement (i.e. top of basement) and the out of sequence thrust. See text for discussion. Seismic section is --13 km long. Location shown in Figures 8 and 13.

796 Allmendinger et al.- Shortening in the Andes at 30S

w E O- Bermejo Basin

2-

8-

Fig. 6. YPF seismic reflection data from the Bermejo Basin (Figure 2), showing the deformed thickness of pre-Mesozoic strata and the unusual depth to the top of inferred basement. Pz- pC(?), deformed Paleozoic(?) and/or Precambrian strata; CO(?), Cambrian-Ordovician(?) limestone. Note that the Cenozoic fold is detached above the reflector at ~7 s (~18 km). Seismic section is about 24 km long.

eastern edge of the thin-skinned thrust belt lies a..t ~15- km depth. This unusually deep dollement and the thickness of the stratigraphic section involved in the thrusting to the west provides important constraints on the amount of shortening in the Precordillera, regard- less of the specific structural geometries chosen. BALANCE OF PRECORDILLERA SHORTENING

Notable features of the Precordillera

The Precordillera presents many challenges to the application of the classical techniques of balanced cross-section construction. Here, we briefly describe several notable features of both the Precordillera and the eastern part of the Iglesia Basin.

Bedding thickness variation and localization of the front of the thrust belt. Substantial Paleozoic tectonic activity severely modified this margin of South America [Baldis et al., 1982; Ramos et al., 1986]. Significant angular unconformities and tightly folded rocks underlie upper Paleozoic and Mesozoic strata in both the morphological eastern Precordillera (here included in the Sierras Pampeanas structural province) and western part of the belt. The eastern two thrust plates of the thin-skinned belt do not show evidence of this strong pre-Cenozoic folding but subtle angular

unconformities do exist and unit thicknesses vary considerably from plate to plate and even within a single plate [Furque, 1979]. This lack of a simple layer cake stratigraphy renders impossible a strict application of Suppe's method [1983; Woodward et al., 1985]. It also negates many of Dahlstrom's [1970] rules, particularly those concerning thrust faults cutting up section in the direction of translation and the juxtaposition of older rocks over younger rocks.

YPF seismic reflection data show that pre-Tertiary units beneath the Bermejo Basin (Figures 2, 6) are deformed and that their present thickness is much greater than that of probably correlative units in the easternmost thrust plate. The change in vergence of Cenozoic structures, from the west-dipping thrusts of the Precordillera to the east-dipping thrusts of the westernmost Sierras Pampeanas, was probably con- trolled by the reversal in taper of the wedge-shaped stratal packages (Figure 7) and by the pre-existing structures beneath the Bermejo Basin. Figure 6 also shows that Andean structures in the foreland basin are detached above a reflector at 7 s (-17 km).

Steep cutup angles. Bedding dips in the Precordil- lera are consistently steep; the shallowest (except for local fold crests and troughs) are ~30 and are located at the eastern side of the thin-skinned belt. In the western part of the belt, this may be due to both pre-


....... base of Tertiary ........... base of Terti.ary ...... ?-":'"' .......... -""-'''' '''''' "'''''' '----- '-" : ---'::.-'-'---' ....... '""'"' '"'" ' ''"'' '"'"... . '-....;4.;. ......................... ........ '- '-- :. . '.-:i" "-' - --'--':-:.':4 ............ -:":'-'' 't;--- ... -':---:--':x-'.-:."?'":v'"":':'"'':':' '"'"'>'::': ........... '"" '""'''":'"..., ,...-':'. ;- -.-"c. .::.-':-.':::::-':-':::::-':::.:.-.:::.:c-- :.

.-:'"t"1-1 I I I I ...!.-.........-:.-...-c.-.-.-,---, " I I ........................ .:.:..::,....,. -. ,.::E.:20 km constrained by seismic reflection data, Figure 5) to the unbroken Cuesta de Huaco anticline in a distance of--20 km (Figure 8). It is possible that the main fault continues northward beneath young basin fill whereas the fold corresponds to the tip line of a minor imbricate. However, the rapid northward loss of displacement may be real, given that many faults display elliptical, rather than linear, displacement functions (Walsh and Watterson, 1989).

Out-of-sequence thrusts. Because of the widespread interest in the self-similar Coulomb wedge model of thrust belts [Davis et al., 1983], out-of- sequence thrust faults have recently received considerable attention. The wedge model requires contin- uous thrusting and thickening throughout, whereas several classical field studies have demonstrated the

cratonward migration of the locus of deformation with time [Geiser and Boyer, 1987; Woodward, 1987; Arm- strong and Oriel, 1965]. Furthermore, Dahlstrom's [1970] "family" of thrust belt structures does not emphasize out-of-sequence thrusts, which violate the basic rules of thrust geometries.

The Precordillera has several prominent examples of out-of-sequence thrusts in both its eastern and western sides. One thrust, particularly well displayed in the field and on seismic data, deformed part of the Niquivil plate after the strata of the plate were rotated to their present dip (Figures 5, 8, and 9). Far- ther west, the Iglesia Basin was filled during movement on the major thrusts within the Precordillera (see tb.e following section) and the basin strata were subse- quently deformed by several well-imaged out-of- sequence thrust faults (Figure 10). The westernmost fault within the Precordillera, the Rfo Las Trancas thrust, overrides westerly derived upper Cenozoic units which onlap (with an angle as large as 40 ) Paleozoic strata within the next thrust plate to the east (Figure 11). This relationship indicates that the thrusts to the east had moved and created topography prior to the deposition of the upper Cenozoic strata. Thus, the Rfo Las Trancas thrust is also a major out-of- sequence thrust; it has a much gentler dip (--25-30 ) than the thrusts to the east and may have as much as 5-10 km of displacement.

Quaternary(?) extensional tectonics. Although not the main focus of this paper, it is interesting to note that Iglesia Basin, as well as its southern extension in the Calingasta Valley, contain a record of young extensional deformation. Whitney and Bastfas [1984] first showed that the north-striking E1 Tigre fault has an easterly dip in shallow trenches; the 50 m high Quaternary scarp is up on the west side, indicating a normal component of displacement. Our studies of arrays of minor faults in natural exposures of the youn- gest Tertiary strata also show dominantly normal faulting with east-northeast horizontal extension


J--" Quaternary alluvium Quaternary alluvial fans

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exposed'j coveredJ faults 0 km 10

hal

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Niquivil

Fig. 8. (Opposite) Geologic map of the eastern part of the Precordillera thrust belt, showing the three easternmost east-verging thrusts in the belt. Based on Furque [1979], Fielding and Jordan [1988], and field observations by the authors. Note the tip line of the Niquivil thrust, southwest of Huaco, and the tip line of the unnamed thrust southwest of Jchal. Seismic line in Figure 5 crosses through the gap in the Cambro-Ordovician block -8 km north of Niquivil. Note the out-of-sequence thrust in the Niquivil thrust plate between Niquivil and Jchal.


..

.

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,., ' ';' "--"'. --". i '.i-",.;.- ";.-.' '"::: ...... '"".i::':'-": :'::' ':;'.:::': :'. , :..-- .-.: :. ....... . --:...'..,:.: .. ....

Fig. 9. Thematic Mapper image showing detail of the anticline associated with the out-of- sequence thrust in the Niquivil thrust plate. Note broad anticline in late Cenozoic strata and underlying unconformity with Paleozoic rocks. Image is 15 km across. See Figure 8 for location.

(Figure 12). The tectonic origin of this episode of young extension is not known. It seems unlikely to be related to development of the high topography [e.g., Dewey, 1988; Burchfiel and Royden, 1985] because, though oriented perpendicular to the topography, it occurs at elevations as low as 1600 m. It appears equally unlikely that it is related to over steepening and spreading of the Precordillera thrust wedge.

Amount of Shortening in the Precordillera

The cross-section in Figure 13 was constructed using a combination of line length and area balancing methods. The hanging wall cutoffs are eroded from every major thrust plate; the gaps between restored thrust plates have been kept to a minimum. This assumption is probably acceptable, at least for several

of the eastern thrusts, because of the proximity of the section line to tip line folds. The section predicts 95 km of shortening of an original 135 km wide belt, or about 70% shortening. Stated another way, the structural thickening has resulted in a tripling of the stratigraphic section.

This unusually large percentage of shortening-- most thrust belts have -50% -- is a direct result of the original thickness of the stratigraphic section and the depth to the dScollement. The constraints imposed by these two features can be investigated by an area balance of the entire deformed package which does not take into account the structures within it. One can visualize this change in area as being accommodated by a series of vertical-sided polygons that get taller and thinner during the deformation. This unrealistic structural geometry gives the absolute minimum


45

W

2 ELEVATION,KM SP700

I 500

I I ! 300

I

LINE 5323 E

100 I !

Fig. 10. A representative seismic line showing the overall geometry of the Iglesia Basin (Figure 2). Station spacing is 50 m. Note the out-of-sequence thrusts on the east side of the basin and lack of evidence for any fault bounding the west side of the basin. The seismic sequences in the Tertiary strata onlap a --12 east-dipping basement surface; they are described in detail in Beer et al. [1990].

amount of shortening; more realistic geometries will give larger values.

Assuming that the original wedge was 4 km thick at its east side and 6 km thick at its west, it must have been 86 km long to match the area of the 37 km wide deformed package (see inset in Figure 13). This estimate is too small because it does not take into account

the area of material eroded from the thrust plates and now residing mostly in the foreland basin (Bermejo basin) to the east. Adding the area of the erosional gaps in the actual restored section (which are mini- mums) and subtracting the area of Tertiary rocks within the thrust plates (Figure 13, bottom), we estimate the pre-erosional area of the deformed package to have been--488 km 2. This area would require that the undeformed wedge was 98 km long, producing 61 km of shortening.

Neither of the resulting values for internal shortening, 49 or 61 km, includes the amount by which the entire package has been translated over the autoch- thon by the easternmost thrust. That horizontal shortening, --23 km, is relatively well constrained by the seismic data. Adding the internal shortening together with the translation gives a total of 72 or 84 km of shortening. Thus, the 95 km of shorteninlg

determined in the detailed cross-section is probably close to the absolute minimum value for the belt. The

actual shortening could be larger although, as we show below, it cannot be too much larger without resulting in an unrealistically thick crust or necessitating tectonic erosion at the plate boundary.

Westward Projection of the Flat dEcollement Balanced cross-sections not only facilitate calcula-

tion of the amount of shortening, but they also enable predictions of how far beneath the mountain belt the d(collement must project. In the present example, if there are no hidden duplexes of basement material beneath the Precordillera, then the d(collement must be flat beneath the Iglesia Basin and extend to a point beneath, or west of the international border. Although, a large amount of seismic reflection data in the Iglesia Basin is available for inspection the two- way travel times are not sufficiently long to image the d(collement at its predicted 16-18 km depth. We cannot rule out the possibility of basement duplexes but a variety of circumstantial evidence, described below, indicates that they may only be present west of the Iglesia basin.


2

'o

no vertical exaggeration

Tertiary Tertiary Ordovician pillow basalt

Ordovician flysch

Fig. 11. Cross-section of the Rio Las Trancas thrust showing the unconformable relations documenting its out-of-sequence nature. Detail views show the onlap of Tertiary strata (in white) on the back of the Rio de Caracol plate and in the Iglesia basin on the back side of the Rio las Trancas thrust plate. See text for discussion.

STRUCTURAL INSIGHTS FROM SEISMIC STRATIGRAPHY OF THE IGLESIA BASIN

The geochronology of the Iglesia basin, in concert with the subsidence history of the Bermejo foreland basin to the east, clearly shows that the Tertiary strata of the Iglesia area accumulated during thrusting in the Precordillera [Johnson et al., 1987; Jordan et al., 1988; Beer and Jordan, 1989; Beer et al., 1990]. The geometry of those stratal packages contains a record of the topographic development on the margins of the Iglesia basin as well as the nature of the structures controlling that development. The seismic stratigraphy of the >350 line kilometers of reflection data collected in the Iglesia Basin has been described in detail in a com- panion paper by Beer et al. [1990]. Here, we briefly review only those aspects which bear on our structural model.

The most striking observation from the seismic data is that, along the western margin, the Tertiary strata onlap onto a basement surface that dips ~12 east (Figure 10). "Basement" here refers to all sub-Tertiary lithologies, which include Carboniferous sedimentary

rocks, Permian granitoids, and Permo-Triassic ignimbrites. Thus, although, there is an important change in topographic slope between the High Cordilleras and the Iglesia basin (e.g., Figure 3), no surface-breaking fault uplifts the former with respect to the latter. The Tertiary strata accumulated against a pre-existing, or synchronously developed, topographic edifice. This fact, the synchronism of the Iglesia and Bermejo basins, and the inferences about the westward projection of the d(collement, strongly suggest that the Iglesia Basin is a "piggyback" basin developed within a single major thrust sheet [Snyder, 1988].

The simplest geometry for producing a piggyback basin is to localize it between two ramps. Ramps to the east in the Precordillera are readily available (e.g., Figure 13). The lack of an exposed thrust on the west side of the Iglesia basin suggests that the uplift of the High Cordillera must be associated with a buried hanging wall ramp; its footwall counterpart must be located still farther west. The observed onlap on the western side of the Iglesia basin is very similar to that described by Medwedeff [1989] for the onlap of syntectonic strata onto the forelimb of a growing fault-bend fold anticline ("growth fault-bend folding").


Equal Area

n 3

N= 57

shortening

C.I. = 2.0 sigma

Equal Area 2

N = 57 C.I. = 2.0 sigma extension

Fig. 12. Equal area lower hemisphere projections of kinematic analyses of minor faults from the Tudcfim region of the Iglesia Valley. All faults collected from strata younger than -7 Ma. Note the dominance of vertical shortening and horizontal, east-west extension. Numbers and black squares show the results of a uniformly weighted moment tensor solution to the fault kinematics (1 is the shortening axis and 3 is the extension axis). Method is described in Marrett and Allmendinger [1990].

:. Tertiary

Paleozoic ............. Cambro-Ordovician Limestone

0 km 20

Iglesia Basin Precordillera Bermejo Basin

Area = 420 sq. kin. Area = 420 sq. kin. Width = 86 kin. Width = 37 kin.

========================================================= ........ ::::-.:::::::.-:: ............................... t ........ '"'""::":':': Generic Area Balance

Fig. 13. Balanced and restored section of the Precordillera. Inset shows at generic area balance of the thrust belt, which is not dependent on the structural geometry within the deformed package and does not include erosion. The scale of the generic area balance is not the same as the structure section. Locations of the seismic lines and profiles in Figures 5, 10, and 11 are shown for reference.


Thus, we suggest that the eastern side of the High Cordillera is the forelimb of a fault-bend fold anticline. The 12 east dip of the basement surface allows us to predict the hanging wall cutoff (13-14 if the dEcollement dips 1-2W) and by implication the footwall cutoff (12-13W) of the ramp responsible for the uplift of the Andes at 30S. These calculations assume that the rules of fault-bend folding [Suppe, 1983] apply to the rocks of the High Cordillera. These assumptions are impossible to prove but seems reasonable in light of the successful application of the tech- nique to basement rocks elsewhere [e.g., Namson and Davis, 1988].

DISTRIBUTION OF SHORTENING IN THE HIGH CORDILLERA

Argentine Side

The Argentine side of the High Cordillera west of the Iglesia basin at the latitude of the cross-section (Figure 2), is characterized by gentle open folding of the Paleozoic strata and apparently minor high-angle faulting (in contrast with observations from farther north [Marin and Nullo, 1988]). Dips in Carboniferous sandstone and quartzite average 12 east-southeast (Figure 14), in rough accordance with the dip of the basement surface seen on the seismic data. The region lacks evidence of significant horizontal shortening (the E1 Indio belt, both Argentine and Chilean parts, is described below), although regional geological studies are not complete enough to rule out the possibility of strike-slip motion along the high-angle faults.

Chilean Side

Three important belts of shortening are present in the High Cordillera: the E1 Indio belt which straddles the international border, faulting in the plutonic nucleus, and the Vicufia fault area to the west. Unfor- tunately, there has been no quantitative assessment of the amount of shortening in any of these belts. In the absence of any more definitive measure, we use a sta- tistical relationship between displacement and fault width (in this case, fault surface trace) to make a rough first order estimate of shortening. Data on 366 faults presented by Walsh and Watterson [1988] shows a log-linear trend of displacement vs. width which can be fit with a straight line with a slope of 1.58. Note that, to convert to horizontal shortening the dips of the faults must also be known.

Because faults in the E1 Indio belt cross the international border, they have been mapped by different workers and their total lengths are difficult to deter- mine. The minimum length of the Barios del Toro fault, ~100 km, indicates 5-10 km of slip. The eastern boundary fault is apparently about half that length and other faults within that belt are much smaller.

Assuming that the fault dips in this belt average about 45 , we estimate about 7-15 km of horizontal shortening. North of 30S, at least 4 major reverse faults occur

Equal Area

o .,Average Bedding: 040 , 12 E

N = 28 C.I. = 2.0 sigma

Fig. 14. Equal area, lower hemisphere projection of poles to bedding in Carboniferous strata from the road up the Agua Negra Pass in the Frontal Cordillera. Note the similarity in dip between the average bedding and the dip of the basement surface seen in the Iglesia basin seismic reflection data (Figure 10).

in the plutonic core of the Chilean High Cordillera [Reutter, 1974; Nasi et al., 1989]. Reutter [1974] estimated shortening at 29S to be on the order of 19 km. This figure is in good general agreement with an estimate based on fault surface trace length, derived from data in Nasi et al. [1989]. Thus, in these two belts, there is probably ~30-40 km of horizontal shortening; this calculation does not include distributed deforma-

tion within the major fault blocks (e.g., folding, pene- trative strain, etc.). Most of this shortening probably occurred prior to 11-16.6 Ma [Maksaev et al., 1984; Kay et al., 1988; Moscoso and Mpodozis, 1988].

The Vicufia fault is 230 km long, which gives between 15 and 30 km slip for a reasonable range of shear moduli; using an average dip of 30 , we estimate 13 to 26 km of horizontal shortening. The amount of this shortening which occurred during Incaic deformation is unknown. The shortening in the Coastal Region is very small, probably no more than 2-3 km. In total, then, the likely amount of horizontal shortening on the Chilean side of the Cordillera is --45-65 km, although the very primitive nature of this calculation must be taken into account.

CRUSTAL-SCALE BALANCE

A whole-crust area balance requires a knowledge of the thickness of the crust prior to the deformation. As neither this figure nor the present day crustal thickness is well known, a virtually unlimited number of models is possible. We make no effort to present all variants below; instead we present one reasonable model which is consistent with the first order boundary conditions described above, but is designed to maximize shortening. Further refinement of the model


must await a more accurate measure of crustal thickness.

Generic Area Balance and Tectonic Erosion/Truncation

Present cross-sectional area. The pin line for these calculations is located at the eastern tip of the Niquivil thrust. This position corresponds to the boundary between the west-dipping, thin-skinned Precor- dillera thrust belt and the east-dipping thick-skinned faults of the westernmost Sierras Pampeanas (Figures 2, 15). Although shortening in the Sierras Pampeanas is important to the overall calculation of crustal shortening, we lack sufficient seismic reflection data to define the fault geometries. Jordan and Allmendinger [1986] estimated shortening in the province at < 5%. However, because most faults in the western Sierras Pampeanas dip to the east, they do not thicken the crust beneath the orogen, which is the main focus of this paper.

The crustal thicknesses are taken to be those presented above. Based on gravity data located south of the transect, it is likely that the present crustal thickness beneath the foreland east of the thrust belt is ~45 km [Introcaso, 1980; Snyder, 1988]. The thickness beneath the crest of the Andes is taken to be 65 km, cor- responding to nearly complete local compensation, and that beneath the Coastal Cordilleras 45 km. These numbers, admittedly poorly constrained, give a cross- sectional area of-17,120 kr 2.

The amount of material eroded from the mountain belt and deposited east of the pin line during defor-

mation must be accounted for. Although the structural geometries in this part of the orogen are relatively two-dimensional, sediment dispersal patterns are in- herently three-dimensional. Lacking any other measure of erosion, we simply use the cross-sectional area of the Bermejo Basin (460 km2), which is well determined because of the excellent seismic coverage (Figures 2, 6, and 15). Adding this value to that above yields a deformed cross-sectional area of 17,580 km 2 for the present width of 380 km (Figure 15, top). This area calculation does not take into account the material eroded from the Chilean side of the orogen but may over-estimate the erosion on the east side because drainage patters show longitudinal transport of material and its concentration in the Bermejo Basin [e.g., Damanti, 1989].

Initial crustal thickness. The upper part of the crust in the Bermejo foreland basin is composed of ~7 km of upper Tertiary strata. These rocks were deposited mostly coincident with the deformation described here (with the possible exception of the Vicufia fault). Thus, assuming a present 45 km crustal thickness and that the Moho has not changed position in the last ~15 Ma, crustal thickness in the foreland prior to the onset of deformation was -38 km. The surface at high elevations (Figure 4) is suggestive of, but does not require, low relief prior to uplift.

No similar reasoning can be made for the internal parts of the orogen, which experienced magmatism prior to ~10 Ma. According to Kay et al. [1987], trace element patterns are compatible with the presence of garnet in the source area of 16.6-11 Ma lavas (the Cerro

;a- 17117 sq. km Width = 383 km

= 458 sq.

%%%%%%%%%%%%

0 50 km 0 ''

50 km

' ...... o:..-'.:::.'.--'...!!!.?.x::s:::::.'.:i:.v;i:-:- :,.! --.':'!:t: i'.- -:st:. :-: '-'--'-:- .... ....._,,............. .. ,, .. ................ , .,. ....

Fig. 15. Generic crustal scale balance of the Andes at 30S. A specific crustal shortening distribution is not assumed. This liberal construction was devised to maximize shortening because it probably over-estimates the present thickness of the crust beneath the High Cordillera, it may underestimate the starting thickness of the the crust, and it neglects any magmatic component prior to ~10 Ma. Total shortening is 137 km or 36%.


de las Tortolas Formation) indicating that the crust had attained significant thickness by the time of their eruption. This age range correlates with shortening due to reverse faults in the plutonic core and E1 Indio belt of the High Cordillera [Moscoso and Mpodozis, 1988] and is coincident with the onset of Precordillera thrusting. Older lavas (the Dofia Ana Formation) show little geochemical evidence of crustal thickening although it is likely that some crustal root was present beneath this older arc [Kay et al., 1987]. The Dofia Ana Formation is geochemically similar to the modern lavas of the Southern Volcanic Zone south of 36S and thus the crustal thickness during the early Miocene at 30S may have been similar to the present, poorly known crustal thickness farther south [Kay et al., 1989].

To maximize the amount of crustal shortening based on the genetic area balance, we show the crust prior to --15 Ma with a relatively uniform thickness across the entire orogen, with the exception of a tapered edge cor- responding to the mid-Miocene plate boundary (Figure 15, bottom). The existence of a mid-Miocene crustal root would reduce the difference in area between initial and final sections and reduce the overall shortening. Thus our starting assumption is probably unrealistically liberal but it does provide one end member possibility.

Calculation and comparison of shortening. Given the geometry shown in the bottom of Figure 15, the starting width must have been 520 km to match the area calculated above. The difference between this

number and the present width of 383 km suggests 137 km of horizontal shortening. This number is, if any- thing, an over-estimate because we have assumed a maximum reasonable value for the present crustal thickness and have neglected any magmatic input even though the ages of the volcanic rocks overlap in time with the age range of at least 30% of the shortening.

The probable value of shortening determined in the previous section of the paper lies between 130 and 170 km. If we compare this observed structural shortening, determined from outcropping structures in the Andes at 30S, it is somewhat larger than what can be accounted for by a reasonable estimation of the present crustal cross-sectional area. However, the observed structural shortening was calculated conservatively and the crustal area balance was a liberal estimate (i.e. our assumptions were designed to maximize shortening in the generic area balance). Thus it is likely that the structural shortening is larger than can be accounted for by the predicted crustal cross-sectional area for the present. Two explanations are likely. First, many authors have suggested that the plate margin has been eroded or truncated, based on the anomalously short distance between the present day trench and the Mesozoic magmatic arc along the coast of north-central Chile [e.g., Rutland, 1971]. Truncation or erosion of crustal material during the Neogene would reduce the area of the present crustal cross-section relative to the amount predicted by surface shortening. A problem with this is the fate of the removed material: if eroded, it might well be underplated because of its buoyancy, and trun-

cation seems unlikely because of lack of field evidence for significant Neogene strike-slip faulting. The other possibility is that the crustal root is much thicker than we have predicted, but the topography is "held down" by the nearly flat subducted Nazca Plate. Given the myriad of unknowns, any attempts at more precise models would be a fruitless exercise.

Distribution of Shortening in the Crust

The surface loci of shortening are well established: the Precordillera accounts for --60-75% of the total

shortening of the upper crust located between the plate boundary and the Sierras Pampeanas craton. The Precordillera also includes nearly all of the post- middle Miocene shortening. In this section, we explore two end member geometries for the link between this upper crustal deformation and the shortening and thickening in the lower crust. We only consider the post 11-16.6 Ma deformation and make no effort to account for older deformation in the High Cordilleras or along the Vicufia fault (as would be necessary in a more complete model). Only the first explains three major observations: (1) the inferred western projection of the Precordillera d(collement, (2) the width of the "lateau" surface on the top of the High Cordillera shown by the hypsometric curve (Figure 4b), and (3) the angle between onlapping upper Tertiary strata and the dipping basement surface at the western margin of the Iglesia basin.

Wedge model. The width of the High Cordillera plateau is about 80 km and the amount of shortening determined for the Precordillera is 95+ km (Figures 4b and 13). Furthermore, we noted above that the Iglesia basin onlap geometry resembled syntectonic sedimen- tation on the forelimb of a major fault bend anticline. Thus, we suggest that the eastern and western slopes of the plateau define the eroded forelimb and backlimb kink-bands, respectively, of what amounts to a major hanging wall anticline (the uneroded forelimb is present only beneath the western Iglesia basin). The angle of the hanging wall ramp, measured from the dip of the basement surface beneath the Iglesia basin is 12 , predicting and 11 footwall ramp. If that ramp is --23 km long and thrusting over the adjacent footwall flat is 95 km, the resulting flat-crested anticline will be 80 km wide and have a structural relief of--5 km, precisely what is observed in the morphology (Figures 4b, 16). In this interpretation, the break in slope along the western side of the High Cordillera is produced by the western kink band arising from the deep footwall ramp and is unrelated to the Vicufia fault. Thus, a critical unresolved problem is the age of last movement of the fault.

This geometry appears to fit the observed morphology and shortening of the Andes at 30S exceptionally well (Figure 16). It does, however, raise a major problem: the ramp which connects upper and lower crustal shortening lies entirely west of the high topography and, by inference, the zone of lower crustal thickening. The plate boundary provides the western limit to the displacement of upper crust with respect to lower crust.


Fig. 16. Crustal wedge model with Iglesia Basin onlap onto forelimb of major anticline. The plateau surface is defined by the profile of maximum elevations (top curve); dotted curve is the average elevation (both at 5x vertical exaggeration) Note that the western slope of the Andes, which coincides with the Vicufia fault, can be explained as a kink band over the footwall ramp.

Thus the displacement transfer to lower crustal shortening and thickening must be via a ramp which is the mirror image of the upper crustal ramp. This geometry (Figure 16) is similar to a crustal-scale triangle zone and has been named "tectonic wedging" by Price [1986]. We have depicted the lower crustal deformation to be geometrically identical to the upper crust for simplicity's sake; shortening could be accomplished completely by crystal plastic mechanisms and be kinematically, though not geometrically, identical.

Duplex model. An alternative means to account for the shortening and thickening is with a duplex zone or zone of distributed ductile thickening beneath the High Cordillera (Figure 17), similar to the model proposed by Isacks [1988] for the Altiplano farther north. It would be difficult to predict the number of horses within the duplex; we arbitrarily assume that they equal the number of major thrusts in the thrust belt. To link lower and upper crustal shortening while using 'easonable ramp angles, the footwall ramp of the frontal horse must be located directly beneath, or even east of, the Iglesia Basin. As in the previous model, the duplex shown could, kinematically, be replaced by

a zone of more homogeneous plastic deformation. Evaluation of models. There are, of course, an infi-

nite range of possibilities lying between these two extremes. Both models account for the uplift of the High Cordillera without a surface breaking thrust on the west side of the Iglesia basin. We lean toward the wedge model because it is simpler to relate to the first order features of the topography and in addition provides a simple explanation for the onlap of Iglesia basin strata onto the titled basement surface. It would be possible, but more difficult, to construct a duplex model which accommodated the topography and the piggy back geometry of the Iglesia basin. In particular, it is difficult to reconcile the necessity of the frontal ramp beneath or east of the Iglesia basin (Figure 17) with the surface geology and available seismic reflection data which indicate a hanging wall ramp-over-footwall flat only west of the basin. The duplex model or some variant works very well where there is a broad region of high topography in the orogen behind the thrust belt. It should be noted that neither model provides an adequate specific explanation for the observed zones of shortening in the Chilean

0 50 km 0

50 krn

'duplex'

%%%%%%%%%%

+++++++

Fig. 17. Crustal duplex end member model. Number of horses arbitrarily chosen to match the number of faults in the Precordillera. This model would be kinematically identical to a distributed ductile shortening model such as that presented by Isacks [1988].


Cordillera. Additional study, especially deep seismic reflection profiles, would undoubtedly show both models presented here to be incorrect to a greater or lesser degree.

CONCLUSIONS

Despite the remaining unknowns, the Andes at 30S represent one of the best opportunities to construct a cross-section through an entire orogenic belt. The surface features are extremely well exposed and generally adequately (though not completely) documented, the plate geometry is known, and the dynamic topography probably reflects first-order crustal structure. Further- more, the lack of magmatism during the main phase of shortening reduces the significance of one of the major unknowns common to most orogens. A major barrier to a complete crustal balanced cross-section is the lack of crustal thickness information. This should constitute a major target for future investigation.

The shortening at the surface in the Precordillera of western Argentina accounts for ~60-75% of the total of ~130-170 km. The 95 km of Precordillera shortening occurred during the last 10-15 m.y. Even though this magnitude is large, the shortening rate, 0.63 - 0.95 cm/yr, is less than 10% of the plate convergence rate. This ratio is in close agreement with other antithetic thrust belts [Allmendinger, 1989]. It is likely that this surface shortening is greater than can be accounted for by present day crustal thicknesses and thus raises the prospect of plate margin erosion or truncation during the Neogene. Alternatively, crustal thickness may be significantly greater than predicted by simple isostatic models; this thickness may not be reflected in the topography due to the influence of flat subduction of the Nazca Plate.

The occurrence of antithetic shortening so close to the trench places some constraints on the number of alternative possibilities for distributing the shortening on a crustal scale. Although not unique, we favor a wedge model in which the main ramp linking upper and lower crustal deformation are located between the plate boundary and the high topography. This model thus differs from that published by Isacks [1988] for the Altiplano of the Andes farther north, where the orogen is much wider. Equally viable is a crustal scale duplex kinematically similar to Isacks' [1988] ductile shortening model. This alternative, however, cannot explain as simply the first order features of the present morphology or the seismic stratigraphy of the Iglesia basin.

Finally, the data presented here bear on models of thrust belt mechanics in both small and large ways. The critical wedge model of Davis and others [1983] requires thickening in the rear of the wedge to maintain its taper. Out-of-sequence thrusts are common in the Precordillera, although whether their shortening is sufficient to maintain a constant taper is unknown [e.g., Woodward, 1987]. The surface slope of the Precordillera (2.25 ) is anomalously low for subaerial thrust belts [Davis et al., 1983]. This might be explained if crystal plastic deformation mechanisms

are dominating or competing with frictional mechanisms along the anomalously deep d6collement. More importantly, however, it appears that major topographic variations can be simply explained by first order ramps and flats in the basal d6collement, a geometry not specifically addressed in the critical taper model. Finally, in the case of the Andes at 30S, the "snow plow" can be nothing other than the plate boundary itself.

Acknowledgments. We are indebted to numerous Argentine, Chilean and North American colleagues for discussions of these issues. Foremost among them are the participants in the Joint YPF-Cornell expedition to the Iglesia Basin and Precordillera in June 1988: L. Alvarez, R. Gorrofio, A. Gutierrez, J. C. Idiart, T. Jordan, E. Kozlowski, and S. Miniti. Further exchanges with V. A. Ramos, T. Jordan, S. Kay have considerably clarified the ideas presented in this report. S.E. Boyer, B.C. Burchfiel, and an anonymous reviewer also provided many helpful comments although they do not necessarily support our conclusions. We are grateful to Yacimientos Petrolfferos Fiscales for permission to publish the seismic reflection data and to the Donors to the Petroleum Research Fund of the American Chemical Society and the National Science Foundation (Grant EAR-8607468) for support. REFERENCES

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(Received June 20, 1989; revised October 31, 1989; accepted October 31, 1989.)

All Men Dinger 1990

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