trench-parallel shortening in the northern chilean forearc...

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ABSTRACT

In the Central Andes, the frictional coupling between South America and the subduct-ing Nazca Plate occurs beneath the Coastal Cordillera of northern Chile. One of the most distinctive characteristics of the Coastal Cor-dillera is a suite of EW topographic scarps located between 19° and 21.6°S latitude. These scarps are associated with predominantly south dipping reverse faults that have almost pure dip slip and produce shortening parallel to the plate boundary. Limited geochronology as well as more regional studies indicate that the scarps formed during the late Miocene and early Pliocene, though some show activity extending into the Quaternary. In several areas, the scarps dammed local drainages, producing internally drained basins that accumulated evaporites. This relationship indicates that the Coastal Cordillera was probably moister during the Late Miocene and Pliocene than it is today and also indicates that the Coastal Cordillera has been signifi cantly uplifted or that the Coastal Escarpment of northern Chile has advanced signifi cantly eastward since the Pliocene. The limited latitudinal extent of the EW scarps and their location symmetrically about the axis of topographic and Wadati-Benioff zone symmetry suggest that they owe their origin to the concave seaward shape of the continental margin due to prior formation of the Bolivian orocline.

Keywords: forearc, Andes, Chile, deforma-tion, geomorphology, tectonics.

INTRODUCTION

Nearly orthogonal plate convergence in the Central Andes (Fig. 1) south of the bend in the coastline at 18.5°S latitude since the early Oligocene has produced a suite of plate bound-ary-parallel tectonic provinces such the forearc (Coastal Cordillera), magmatic arc, and back arc (Subandean belt). In the Coastal Cordillera of northern Chile between 19° and 21.6°S latitude (Fig. 2), however, the most prominent morpho-structural features are fault scarps orthogonal to the plate boundary and approximately parallel to the convergence direction. Because these fea-tures are restricted to the Coastal Cordillera (the only part of subaerial South American crust in direct contact with the Nazca Plate) but are also latitudinally restricted to just part of northern Chile, they must contain information about the nature of three-dimensional plate coupling.

These scarps have previously been inter-preted as the product of domino-style normal faulting (e.g., Reijs and McClay, 1998), but we show here that they are associated with moder-ately dipping reverse faults. With no signifi cant oblique slip component, these faults produce horizontal shortening almost exactly parallel to the trend of the Coastal Cordillera and the plate boundary. In this paper we document the structural geometry, kinematics, and ages of these structures. Possible explanations for the tectonic origin of the structures must take into account their limited geographic distribution. We conclude that they are fundamentally related to the concave-seaward shape of the Bolivian orocline, though they are probably not a product of bending itself.

The margin-orthogonal scarps also control the evolution of drainage systems in the Coastal

Cordillera. The history of river incision and abandonment related to these features refl ects three fundamental processes, one climatic and two tectonic: (1) a climatic transition from semiarid to hyperarid environment after the Miocene, (2) eastward base-level migration due to coastal escarpment retreat, probably con-trolled by subduction erosion, and (3) Coastal Cordillera uplift due probably to underplating in or near the interplate seismic zone.

FOREARC TECTONIC SETTING

Plate Setting and Lithospheric Structure

The Nazca-South America plate boundary between 18 and 24°S latitude has long been considered most typical of an “Andean mar-gin.” The subducted plate beneath northern Chile (Fig. 1) dips ~30° eastward beneath the continent (Cahill and Isacks, 1992). More than a decade of intense geophysical exploration has provided detailed images of the subducted plate and its interface with South America (ANCORP Working Group, 1999; Buske et al., 2002; Götze et al., 1994; Husen et al., 2000; Husen et al., 1999; Wigger et al., 1994; Yuan et al., 2002; Yuan et al., 2000). The Peru-Chile Trench is located just 70–150 km offshore (von Huene et al., 1999). The Nazca-South America con-vergent plate boundary is responsible for some of the largest earthquakes in recorded history (Comte and Pardo, 1991; Tichelaar and Ruff, 1991). The 1995 Mw = 8.1 Antofagasta inter-plate earthquake, which occurred just south of the region described here, showed that the inter-plate seismic zone extends from 20 to ~50 km depth (Delouis et al., 1997; Husen et al., 2000; Husen et al., 1999; Pritchard et al., 2002).

Trench-parallel shortening in the Northern Chilean Forearc: Tectonic and climatic implications

Richard W. Allmendinger†

Department of Earth and Atmospheric Sciences, Cornell University, Ithaca, New York 91125, USA

Gabriel GonzálezDipartamento de Ciencias Geológicas, Universidad Católica del Norte, Antofagasta, Chile

Jennifer YuGreg HokeBryan IsacksDepartment of Earth and Atmospheric Sciences, Cornell University, Ithaca, New York 91125, USA

GSA Bulletin; January/February 2005; v. 117; no. 1/2; p. 89–104; doi: 10.1130/B25505.1; 15 fi gures; Data Repository item 2005022.

†E-mail: [email protected].

ALLMENDINGER et al.

90 Geological Society of America Bulletin, January/February 2005

Geologic Setting

The Coastal Cordillera of northern Chile is formed mainly by Jurassic-Early Cretaceous dio-ritic to granodioritic plutons and Jurassic volca-nic rocks. These units comprise the remnants of a Mesozoic magmatic arc formed at the birth of the modern Andes (Coira et al., 1982; Mpodozis and Ramos, 1990). The magmatic arc was emplaced on an ensialic crust composed of Paleozoic sedi-

mentary and Precambrian metamorphic rocks. The most important structure of the Coastal Cordillera is the Atacama Fault System (Arabasz, 1971; Brown et al., 1993; Riquelme et al., 2003; Scheuber and González, 1999) that extends for more than 1000 km between the 21–26ºS lati-tude. Within our study area, the extreme northern segment of the Atacama Fault System has been reactivated during the late Cenozoic as the Salar Grande fault (Fig. 3).

The Neogene to Quaternary sedimentary record of the Coastal Cordillera attests to pre-dominating arid and hyperarid climatic condi-tions. Several internal basins in the Coastal Cordillera are composed of Oligocene-Miocene alluvial deposits covered locally by Mio-Plio-cene evaporite deposits (Chong et al., 1999; Hartley and Chong, 2002; Hartley and Jolley, 1995), including the Salar Grande described in this paper (Fig. 3).

Figure 1. Regional morphology of the Nazca-South America plate boundary between 10° and 35°S latitude. Onshore digital topography from the USGS 30 arc second DEM; marine bathymetry from the ETOPO 5 DEM. The line labeled “Gephart Symmetry Plane” is the plane of bilateral symmetry from Gephart (1994). Wadati-Benioff zone contour in kilometers depth from (Cahill and Isacks, 1992). Box shows the location of Figure 2. The approximate rupture extent of large historic earthquakes (shaded ovals) is after Pritchard et al. (2002).

FOREARC DEFORMATION, NORTHERN CHILE

Geological Society of America Bulletin, January/February 2005 91

Figure 2. Shaded relief map of northern Chile based on the new Shuttle Radar Topographic Mapping Mission 90 m DEM of northern Chile. Boxes show the locations of the study areas detailed in subsequent fi gures. Box labeled “BA” is Barranco Alto. Camarones marks the Quebrada de Camarones. Conical mountains in the eastern part of the DEM along the Chile-Bolivia international boundary are volcanoes in the Western Cordillera, which marks the western boundary of the Altiplano. The solid and dashed white line shows the location of the topographic transect used to calculate the magnitude of the plate boundary parallel shortening; the solid segment of the transect is shown in Figure 12.

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Topography and Late Cenozoic Tectonics

The Bolivian Orocline has long been recog-nized and was one of the fi rst major tectonic features to be studied using digital elevation models (DEMs) in the late 1980s. Isacks (1988) compiled and used a DEM to evaluate the amount of shortening and rotation along strike in the Central Andes. He ascribed the orocline to the fact that shortening is greatest at the bend in the Andes at Santa Cruz, Bolivia, and dimin-ishes to the north and south. Gephart (1994) used Isacks’ DEM to show that the topography of the central Andes has a well-developed bilat-eral symmetry (Fig. 1) with the symmetry axis defi ned by a great circle whose pole coincides with the pole of rotation between Nazca and South America during a relatively stable period

from 36 to 20 Ma at the start of modern Andean mountain building. Gephart’s symmetry plane crosses the coastline of northern Chile at 20.5°S latitude (Fig. 1), south of the bend in the coast-line at Arica, Chile (18.5°S).

The Coastal Cordillera forms a 1000–1800 m structural high (Fig. 2), bounded on its western side by a 1000 m coastal escarpment (Paskoff, 1980). North of Iquique the escarpment extends down to and is currently being undercut by the ocean. In contrast, south of Iquique the escarp-ment is inactive and fl oored by 1000–4000 m wide coastal plains that contain Late Pleistocene marine terraces (Radtke, 1987; C. Casanova, per-sonal commun., 2003). The most prevalent Late Cenozoic structures in the Coastal Cordillera as well as offshore are trench-parallel normal faults (Niemeyer et al., 1996; von Huene and Ranero,

2003; von Huene et al., 1999). Although some have suggested that the Coastal escarpment is a single normal fault scarp (Armijo and Thiele, 1990), there are no outcrops of this inferred fault and many of the exposed normal faults dip in the opposite direction toward the Andes (Delouis et al., 1998; González et al., 2003). The eastern side of the Coastal Cordillera is formed by gentle topography with an irregular mountain front embayed by the sedimentary infi ll of the Central Valley.

Our study of the forearc of northern Chile employs a new high-resolution DEM produced with radar interferometry (InSAR). These data cover all of northern Chile at a 20 m resolution, providing detailed images of Late Cenozoic fault scarps and other features of tectonic and geo-morphic interest. One of the fi rst-order regional

Figure 3. Shaded relief map of Salar Grande and the Chuculay system of EW-striking fault scarps. Sites A–D are described in the text. Stereographic lower hemisphere equal area projections show the available fault slip data and have been shaded as fault plane solutions (the “T” quadrant in gray and the “P” quadrant in white). The principal axes of infi nitesimal strain from fault slip analysis are labeled “1” for extension and “3” for shortening. The data, though sparse, consistently indicate that the EW faults are reverse faults.



observations made from these data are that the Salar Grande segment (Fig. 3) of the Atacama fault zone and related structures are character-ized by a much more “youthful” appearance (sharper, better defi ned fault scarps) than regions farther south (Yu and Isacks, 1999).

EAST-WEST REVERSE FAULTING

Regional Distribution and Character

The prominent E- to ENE-striking fault scarps are restricted in latitudinal extent from just south of the Río Loa at 21.6°S to the Quebrada de Camarones at 19°S (Fig. 2). The scarps are as much as 450 m high and 35 km long. The associated faults are completely restricted to the Coastal Cordillera, becom-ing blind and plunging beneath the strata of the Central Valley. The majority of the scarps face northward, although between Iquique and Salar Grande there are several well-developed south facing scarps. The backsides of many of the scarps display a gently inclined ero-sional surface on the Mesozoic “basement” of the Coastal Cordillera. At Iquique, there is a change in the orientation of the scarps: to the south the scarps strike predominantly E or ESE, but to the north they strike consistently ENE. This change in scarp orientation accompanies a change in regional trend of the Coastal Cordil-lera from north-south to the south or Iquique to NNW-SSE to the north. Thus the scarps are everywhere approximately perpendicular to the trend of the Coastal Cordillera.

Key Field Localities

In a hyperarid environment, there is no surface water to carry away the products of mechanical weathering and the topography sim-ply becomes buried in its own debris. The lack of fault exposure and open fractures on the faces of the scarps, combined with characteristic tilted blocks led most previous authors to interpret the structures as domino normal fault blocks (e.g., González et al., 1997; Reijs and McClay, 1998). However, some of the structures are exposed where they cut the Coastal Escarpment and there are a few scattered exposures in the deep canyons north of Pisagua. In every case, the fault plane dip, the associated scarp morphol-ogy, and kinematic analysis of fault plane stria-tions demonstrate reverse fault motion during the Late Cenozoic.

ChuculayThe Chuculay system, mapped by Skarmeta

and Marinovic (1981), comprises several EW scarps located east and south of Salar Grande

(Figs. 2 and 3). Though not the main topic of this paper, one of the most remarkable structures in the region is the Salar Grande fault scarp, which produces ~10 m of vertical offset of the halite surface of the Salar Grande (“A” in Fig. 3). To the north of the region shown in Figure 3, the Salar Grande fault produces several tens of meters of horizontal displacement of drainages, suggest-ing that it is predominantly a strike-slip fault (González et al., 2003). The Salar Grande fault is not offset and is therefore younger than the suite of EW scarps of the Chuculay system. The main scarp at Cerro Chuculay is ~350 m high and it and several related faults have produced a suite of well-defi ned blocks tilted ~4–5°S (“B” in Fig. 3). Two strands of the Chuculay fault system offset a small antecedent drainage (“C” in Fig. 3) that is no longer active. The northern strand truncates and drops down the catchment area (now a small closed basin) relative to the channel, and the southern strand offsets the same channel vertically but not horizontally, indicat-ing no strike-slip motion on the fault.

Two outcrops in this system yielded expo-sures of the bounding faults. To the south along the Coastal Escarpment, a brittle gouge zone in Mesozoic igneous rocks dips to the south beneath the uplifted block of the scarp (“D” in Fig. 3). The fault is thus a reverse fault. Like-wise, a small exposure of the fault plane along the scarp located just north of Cerro Chuculay shows reverse displacement with nearly down-dip slickensides. The age(s) of these structures is (are) not known. They produce a broad smooth scarp in the surface of the Salar Grande, however, suggesting at least some movement post-dating salar halite deposition.

Barranco AltoAlthough one of the smaller features in the

region, the 60 m high, south-facing scarp at Barranco Alto (Figs. 2 and 4) is particularly important for its well-preserved age relations. The structure consists of two scarps: a northern, more continuous scarp and a parallel southern scarp that dies out ~0.5 km inland from the Coastal Escarpment. The former scarp yielded no outcrop but the southern one (“A” in Fig. 4) highlights some key relationships. A small evap-orite basin occupies the depression immediately south of the southern scarp, and outcrops on the edge of the Coastal Escarpment show that the strata of this basin onlap the fault scarp (Fig. 5). The reverse fault that produced the scarp also cuts into the basin strata and produces a subsid-iary 8–10 m scarp.

We were able to date a reworked tuff in the basin via single crystal laser fusion ages on 19 feldspar crystals, predominantly plagioclase. Eliminating two obvious xenocrysts, the 17

remaining crystals yield a statistically signifi cant isochron (Fig. 6)1 that accounts for the excess argon in this sample. Based on the isochron, the best estimate for age of the tuff is 5.62 ± 0.1 Ma (T. Spell, personal commun., 2002). Large euhe-dral biotite crystals in the sample indicate that the reworking is minor and that the age obtained is probably close to the age of accumulation of the surrounding strata in the basin. The tuff is par-ticularly signifi cant as it is part of the section that onlaps the scarp in the upper plate of the reverse fault and is itself offset by the fault where it cuts the basin. Though the total separation of the base-ment-basin contact is ~60 m, the tuff is offset by just 2–3 m in a thrust sense. Thus, the Barranco Alto site shows that at least the southern of the two EW scarps accrued most of their displace-ment prior to 5.6 Ma, but also experienced minor thrust motion after 5.6 Ma (Fig. 5).

The topography of the Barranco Alto area records a relationship between scarp evolution and paleodrainages that we see repeated at many sites. A linear drainage (“B” in Fig. 4) cuts straight across the uplifted hanging wall near the east end of the main scarp. This small valley once provided continuous drainage from the interior of the Coastal Cordillera toward the sea although it is now folded so that the north-west segment slopes toward the coastal escarp-ment and the southeast segment slopes toward the eastern end of the Barranco Alto evaporite basin. Furthermore, there is no sign of the valley to the southeast of the scarp, suggesting that the basin covers the old paleovalley. Thus, the mor-phology records the following history (Fig. 7).

Sometime prior to 5.6 Ma, there was suf-fi cient moisture in the Coastal Cordillera and/or the Central Valley to cut a small valley that drained from the interior to the ocean. In the Pliocene, the Barranco Alto east-northeast trending reverse fault began to uplift. Initially, the stream had enough power to incise the uplifting block, cutting the well-defi ned chan-nel we see today. The continued uplifting and folding of the hanging wall eventually defeated the drainage and a closed evaporite basin formed in the footwall. Strata accumulated in the basin, onlapped the growing fault scarp and covered the paleovalley in the footwall. If the coastal escarpment were located where it is today, the growth of the fault scarp would have only diverted the drainage to run along the scarp to the southwest until it reached the escarpment. Instead, higher topography must

1GSA Data Repository item 2005022, tables con-taining the original analytical data for the geochronol-ogy displayed in Figures 6 and 9, is available on the Web at http://www.geosociety.org/pubs/ft2005.htm. Requests may also be sent to [email protected].

ALLMENDINGER et al.


have once existed farther west to block the outfl ow. The coastal escarpment, if it existed at 5.6 Ma, must have been located still farther west. This conclusion and the fact that the paleovalley itself is truncated at the escarpment (“B” in Fig. 4) indicate escarpment retreat since 5.6 Ma. Defeat of the channel described here must have been a function of both uplift rate of

the hanging wall and diminishing water supply (roughly correlated with precipitation) (Sobel et al., 2003), and in fact the subsequent fi lling of the evaporite basin indicates that the upstream portion of the drainage continued to have suf-fi cient water power, at least intermittently, to carry medium grained sediment. There must have been more water in the Coastal Cordillera

when the channel was cut prior to 5.6 Ma than there is today, as there is insuffi cient water in the Coastal Cordillera under current climatic conditions to cut an incised valley, and virtually all evaporite basins there are fossil.

Although the relationships are less clear, the scarp south of Barranco Alto (“D” in Fig. 4) also has a sandy, evaporite basin only in the footwall

Figure 4. Shaded relief map of the Barranco Alto system of ENE-striking fault scarps. Sites A–D are described in the text. Lower hemisphere equal area projections are plotted as described in Figure 3. The box at site “A” shows the location of the detailed geologic map in Figure 5.



of the reverse fault that formed the scarp. These detailed observations are in general accord with the post ca. 6 Ma drying out of northern Chile suggested by Hartley and Chong (2002), although nothing in our data constrains whether the onset of hyperaridity was at 6 or signifi -cantly earlier than 6 Ma.

The Barranco Alto site is interesting for another reason: If one projects the northern fault scarp westward across the Coastal Escarpment, it coincides with a step in the coastal marine ter-race of ~40 m (Fig. 4, “C”). Although no young fault zone was found in the coastal exposures (perhaps because of cover due to modern beach and sand dune deposits), the facing of the step in the coastal terrace is the same as that of the scarp. The paleontological record on the coastal marine terraces indicates that they are Pleistocene in age (Casanova, personal commun., 2003).

PisaguaTo the north of Iquique, the EW scarps are

sparser than to the south, but are much lon-ger and more distinct. Furthermore, the coast

north of Iquique lacks the narrow but relatively continuous marine terrace that characterizes the coast to the south. Instead, the coastal escarp-ment plunges directly into the ocean. The Pisagua scarp (Figs. 2 and 8), ~65 km north of Iquique, is a prominent north-facing scarp ~25 km long and 160–260 m high. The scarp can be traced from the coastal escarpment to the western edge of the Central Valley. The geology of this area was mapped by Silva (1977). We describe this structure below, from west to east.

Similar to, but better developed than at Barranco Alto, the Pisagua scarp is associated with a narrow uplifted marine terrace that is not present either to the north or south along the coast. Exposures of the fault at the escarp-ment show it to be an east-northeast trending reverse fault that dips to the south (locality “A,” Fig. 8). A west-dipping, north-striking normal fault with ~400 m of vertical throw is well exposed in the Quebrada Tiliviche just north of the town of Pisagua (Fig. 8). The trace of the Pisagua reverse fault is offset by the normal fault at site “B,” suggesting as we have

seen elsewhere, that the latter is the younger structure. A similar relationship between a smaller north-striking normal fault and the Pisagua fault scarp may be exposed at “C,” but the throw on the normal fault is considerably less and therefore the offset of the Pisagua fault scarp is more open to question.

Somewhere between sites “D” and “E” in Figure 8, the Pisagua scarp changes from a fault scarp (to the west) to a fold scarp (to the east). At site “E,” the fault is blind and instead, a well-developed tip line fold is exposed. The scarp is much smoother and has a lower slope angle and less relief than farther west. The valley that cuts and incises the scarp at “E” has no catchment area but simply disappears beneath the fl at surface of the Central Valley at site “F.” Thus, we conclude that, like the example at Barranco Alto, this is a paleovalley that predates and is synchronous with the initial ENE reverse fault-ing but has been largely abandoned and covered by the youngest deposits of the Central Valley.

A volcanic tuff is folded above the tip line along with the rest of the sedimentary section

Figure 5. (A) Simplifi ed geologic map of the Barranco Alto area and (B, C) cross section showing the relations between faulting, depositional onlap of the fault scarp, and the dated tuff horizon.

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Figure 6. Isochron plot of the dated tuff from Barranco Alto collected at the location shown in Figure 5. Analysis performed by the Nevada Isotope Geochronology Labora-tory and interpreted by T. Spell (personal commun., 2003).

Figure 7. Schematic interpretation of the evolution of geomorphic and structural features at Barranco Alto. See text for discussion.



at site “E.” The sample of this tuff collected for isotopic dating yielded a separate of very thin, fi ne-grained biotites with as much as 20%–30% plagioclase. Because these are volcanic rocks that cooled very rapidly, the mixture of the two minerals should not produce a signifi cant prob-lem. Stepwise heating yielded a total gas age of 6.36 ± 0.03 Ma, but did not produce a plateau. However, an isochron for steps 5–10 produced a reliable age of 3.49 ± 0.04 Ma (T. Spell, per-sonal commun., 2004), which we consider to be the most reliable age for this sample (Fig. 9). These data demonstrate that the Pisagua struc-ture was active into the late Pliocene, further lending credence to the suggestion above that the uplifted Pleistocene marine terrace at the coast is a produce of fault movement.

Just east of site “E” the Pisagua structure bifurcates into two tipline folds, one trending NE and the other nearly EW. The fold and the tipline of the fault are well exposed in the bottom of the Quebrada Tiliviche at site “G” (Fig. 8). Here, the Tiliviche drainage, which is active today, has been superimposed on and

downcut through the Mesozoic igneous base-ment in the core of the Pisagua fold. Fault exposures show that the fault continues to be a reverse fault; the fanning of shear planes and the geometry of the tip line fold suggest that it may be a trishear-like structure. East of the Quebrada Tiliviche, the Pisagua structure has virtually no surface expression at all.

AtajañaThe Atajaña scarp, located just south of the

Quebrada Camarones (Fig. 2), is the longest and highest, as well as the northernmost, of all the east to east-northeast striking scarps. North of the Quebrada Camarones the Coastal Cordillera decreases in elevation and trends out to sea; at Arica, the Coastal Cordillera is completely sub-merged. It is possible that east to northeast-strik-ing reverse fault scarps continue to be present in the submerged part of the Coastal Cordillera north of Atajaña.

Like most of the structures of the region, the fault that produced the Atajaña scarp is not gen-erally exposed, but outcrops in the Quebrada de

Chiza (“A” in Fig. 10) and along the coast (“B,” Fig. 10) provide critical insight. Along the Pan-American Highway where it descends into the Quebrada Chiza, a nucleus of Mesozoic igneous rocks and folded Cenozoic strata cut by minor faults is exposed. Exposures in the bottom of the canyon show that the Atajaña structure is a blind thrust with an overlying fault-propagation fold. The folded layers diminish in dip up-section and the fold wavelength becomes much broader; the entire structure resembles a trishear fault-propa-gation fold (e.g., Allmendinger, 1998). Exactly where the tipline pierces the surface west of site “A” is unknown. A good candidate would be somewhere near point “C” where the scarp becomes gentler, lower, and smoother.

Along its entire strike length, the hanging wall of the Atajaña fault is composed of Creta-ceous red beds whereas the footwall is made up of Jurassic volcanic rocks (Anonymous, 2003). This relationship suggests that the fault may have originated as a Mesozoic normal fault and was subsequently, at a much later time, reacti-vated as a Cenozoic reverse fault.

Figure 8. Shaded relief map of the Pisagua fault scarp. The lower resolution topography west of 70°10′W is from the Shuttle Radar Topo-graphic Mapping Mission (SRTM) 90 m DEM for northern Chile as our 20 m DEM does not cover the Pisagua peninsula. Sites A–G are described in the text. Lower hemisphere equal area projections are plotted as described in Figure 3.

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Figure 9. Isochron plot from volcanic tuff folded above the tipline of the Pisagua thrust, from site E along the Pisagua fault scarp (Fig. 8). See text for discussion. Analysis per-formed by the Nevada Isotope Geochronol-ogy Laboratory and interpreted by T. Spell (personal commun., 2003).

Figure 10. Shaded relief map of the Atajaña scarp, highlighted by the arrows. Sites A–C are described in the text. Box shows location of detailed topography displayed in Figure 11. Note the spatial association of the fault and the warped marine terrace fragment exposed at the base of the Coastal Escarpment.



As at Pisagua and Barranco Alto, an upwarped Pleistocene coastal terrace remnant is present where the Atajaña scarp intersects the Coastal escarpment (Fig. 10, site “B”). A detail of the topography (Fig. 10, inset) shows that the terrace is not offset but is present at the same elevation both north and south of the projected trace of the fault. This suggests localized upwarping, though small scarps beneath the resolution of the topog-raphy are possible.

REGIONAL SYNTHESIS

Orientation and Magnitude of North-South Shortening

Though outcrops are few and exposures of fault planes scarcer still, we were able to col-lect fault slip measurements on just over 20 fault planes, all from the fault zones of the major EW striking reverse faults. Following the P and T axis method of Marrett and Allmendinger (1990), we calculate the infi nitesimal strain axes (the so-called P and T axes) and plot them in Figure 11. The average shortening axis (“3” in Fig. 11) is nearly horizontal and trends ~170° and the exten-sion axis is nearly vertical. The subhorizontal intermediate principal axis (“2”) indicates that, on average, the deformation has almost no strike-slip component. The shortening azimuth is very nearly exactly parallel to the regional trend of the Coastal Cordillera in this part of northern Chile.

In this hyperarid region, which has experi-enced little erosion since the Pliocene and Late Miocene when the faults were mostly active, scarp height is an excellent proxy for the verti-cal throw on the fault. Scarp heights were deter-mined on a topographic profi le parallel to the Coastal Cordillera and crossing all of the major fault scarps (white line in Fig. 2). The southern third of this profi le, shown in Figure 12, shows the asymmetry of the scarps and emphasizes not only the height but also the tilting of the blocks. In general, there is a rough correlation between the degree of tilting and the scarp height (or composite scarp height), indicating that tilting and fault slip are related. In fact, removing the effect of the tilting would make the average shortening axis exactly horizontal.

Assuming that the faults associated with the scarp dip 45°, which fi ts with what is known of the regional kinematics (Fig. 11), the mag-nitude of the horizontal shortening will be about the same as the vertical uplift. If so, the EW reverse faults produce, in total, a little more than 3 km of horizontal shortening in a distance of ~300 km, or ~1% shortening. If the faults average 60° dips, then the shortening is ~1.8 km across the same distance. The percent shortening is marginally greater in the southern

Figure 11. Summary diagram of the fault kinematics for all of the measured EW reverse faults, displayed on an equal area, lower hemisphere projection. The data are displayed as described in Marrett and Allmendinger (1990), with the solid circles representing “P” axes and the open boxes “T” axes. The statistical best fi t principal infi nitesimal strain axes labeled “1” (extension), “2,” and “3” (shortening) were used to derive a pseudofault plane solution, shown gray.

Figure 12. Topographic profi le constructed parallel to the Coastal Cordillera. Location of profi le is shown as dashed white line in Figure 2. Numbers on vertical lines show the vertical relief of each fault scarp and the numbers in degrees show the angle of tilt of the mid-Ter-tiary erosional surface.

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end of the transect than in the north, which is not surprising given the poor development of scarps in the region between Iquique and Pisagua.

Relationship to Coastal Cordillera Drainages

At several localities in the Coastal Cordillera small, narrow, generally straight and shallow valleys lie totally abandoned, with no evidence that they have carried water in a very long time. These valleys generally lack a catchment area and incise the EW scarps, but where they inter-sect deep canyons or the Coastal Escarpment they are left truncated and hanging, with no evi-dence that they even attempted to cut down to modern base level. The small valley at Barranco Alto has already been discussed in some detail above. Another example occurs just south of the Río Loa (Fig. 13). There, the southernmost out-crop of the EW fault scarps has been cut by two small channels. They are inactive now and their merged channel hangs more than 600 m above the bottom of the Río Loa canyon.

In all of these cases, there was enough water in the Coastal Cordillera when the EW scarps began to form, so that these small drainages could begin to incise the scarps. In many cases, the footwall blocks of the EW scarps have small evaporite basins that show no evidence of ever having been present in the hanging wall. Suffi cient moisture continued in the area even after the uplift of the EW scarps outpaced the ability of the streams to continue downcut-ting. The scarps formed dams that impounded the drainages and produced internally drained basins. We know from our data at Barranco Alto, as well as more regional data on the ages of evaporite deposition in the area (Chong et al., 1999; González et al., 1997; Hartley and Chong, 2002), that this wetter environment persisted until at least 5.6 Ma and if the age of the Pisagua tuff is reliable, until at least ca. 4 Ma.

The truncation of the small valleys shown in Figure 13 indicates more than 600 m of down-cutting of the Río Loa since those valleys were active. In light of the proximity of this area to

the current coastline, this downcutting could have been accomplished in either of two ways (Fig. 14). One possibility is that most of the uplift of the Coastal Cordillera is mid-Miocene and younger. Alternatively, the Coastal Escarp-ment could have retreated signifi cantly eastward since the drainages became inactive at the begin-ning of the Pliocene. Again, the relations at Bar-ranco Alto require some eastward retreat of the escarpment but we cannot say how much.

Either possibility has important though dif-ferent tectonic implications and both were probably active simultaneously. As there is no signifi cant continental shelf or beveled marine platform offshore, it is likely that any signifi cant Coastal Escarpment retreat must be the result of tectonic (i.e., subduction) erosion of the forearc. This process has been shown to have operated along this continental margin, with a very rough average rate of erosion and retreat of 1 km/1 Ma (von Huene and Ranero, 2003; von Huene and Scholl, 1991; von Huene et al., 1999). Although fl exure of the locked plate boundary during the

Figure 13. Shaded relief map of the southernmost EW fault scarp just south of the Río Loa, showing the relationship between drainage development, EW faulting, and canyon incision. Sites A–C described in the text.



interseismic part of the earthquake cycle can explain transient uplift of the Coastal Cordillera, the most popular explanation of long-term uplift is the underplating of material to the base of the continental crust beneath the Coastal Cordillera (Armijo and Thiele, 1990; Delouis et al., 1998). In this process, material eroded from the leading edge of South America is transported downdip to at or near the base of the interplate seismic (~50 km) where it is reincorporated at the base of the South American crust. The resultant thickening of the continental crust produces isostactic uplift.

At the present time we cannot distinguish between the two. Further insight into the history and mechanisms of river incision in the Coastal Cordillera could provide our most exquisite record yet of subduction erosion and underplat-ing during the last 10 m.y. of geologic history.

TECTONIC INTERPRETATION

Any interpretation of the EW faults must take into account three basic facts. (1) They are limited to the Coastal Cordillera, the only part of South American continental crust in

direct contact with the subducted Nazca Plate. As such, they must in some way be related to the coupling between the two plates. (2) Their shortening direction parallels both the Coastal Cordillera and the contours on the underlying Wadati-Benioff zone (Fig. 1), further suggest-ing a link to plate coupling. And, (3) they are limited in geographical extent to the region between 19° and 21.6°S latitude. Thus, there must be something unique about the area in which they occur and they do not refl ect gen-eral plate coupling processes.

The Shape of the Bolivian Orocline

The EW reverse faults are not spatially related to the bend in the coastline at 18.5°S but occur entirely to the south. However, the symmetry plane that best defi nes the Bolivian Orocline (Gephart, 1994) does not coincide with the bend in the coastline either, but crosses the coastline nearly 2° of latitude farther south at 20.5°S (Fig. 1), just south of the city of Iquique. The EW faults are distributed equal distances to the north and south of Gephart’s symmetry plane. This striking spatial coincidence leads

us to the conclusion that the faults are related to processes occurring on the inner arc of the Bolivian Orocline.

There are two ways in which oroclinal bending might infl uence the deformation in the Coastal Cordillera. First, the buckling of a beam produces beam-parallel shortening on the concave side of the neutral surface. Thus, conceivably the Coastal Cordillera, which is located on the concave side of the Bolivian Orocline, has experienced minor shortening parallel to the coast. While appeal-ingly simple and easy to understand, we think it unlikely that oroclinal bending is mechanically so simple. Furthermore, recent paleomagnetic data suggest little or no vertical axis rotation of the forearc during the time that the EW structures were active (e.g., Roperch et al., 1999).

Alternatively, it may be that the preexisting curved shape of the plate boundary induces a deformation fi eld with a component of compres-sion parallel to the boundary. Bevis et al. (2001) have modeled the interseismic velocity fi eld of an elastically deforming forearc of a locked plate boundary with concave curvature (Fig. 15A). In the region overlying the zone of interplate cou-pling and locking, the velocity vectors converge

Figure 14. Diagram illustrating the two mechanisms by which the Río Loa could have been incised. Dark profi le in each case represents the modern setting wheras lighter profi les represent older positions of the Coastal Cordillera and Escarpment. The paleodrainage shown in the two scenarios represents the intersection at site B (Fig. 13) of the small paleodrainge south of the Río Loa canyon with the Río Loa.

ALLMENDINGER et al.


strongly toward the symmetry plane. One can calculate two-dimensional strain for any three non-colinear points knowing their initial and fi nal positions (or initial positions and displacement vectors). In the case of the model of Bevis et al., both principal axes of strain are negative (i.e., a non-area constant deformation), and the shorten-ing perpendicular to the symmetry plane is larger than that parallel to the convergence direction.

The forearc of northern Chile is signifi cantly less curved than that shown in the model of Bevis et al., and the GPS stations are much more sparsely distributed (Fig. 15B). The triangle of

GPS stations that directly overlays the Gephart symmetry plane (IQQE-PPST-PTCH) does, in fact, yield two negative principal axes, indicating shortening in all directions. However, the most negative axis is oriented at a low angle to the EW scarps and oblique to the convergence direc-tion. Furthermore all of the other triangles yield essentially EW shortening and NS extension. Two dimensional analysis of just the coastal GPS stations suggests that, in general, there is minor shortening between those stations, particularly if the three northern stations (ARIC-PSAG-VRDS) are analyzed separately from the three southern

stations (VRDS-PTCH-PBLN). Thus, the cur-rent GPS data displays some orogen parallel shortening in the region of the symmetry plane, consistent with the Bevis et al. hypothesis, but the case is less than compelling. Additionally, there is no guarantee that the interseismic elastic strain is converted into permanent strain in any straight-forward manner.

Our very limited geochronology data suggest that most of the shortening occurred during or prior to the early Pliocene, though there may be minor reactivation during the Quaternary. As recently summarized by Kendrick et al. (2003),

Figure 15. (A) Model of the velocity vectors during interseismic deformation in the upper plate of a locked curved plate boundary (modifi ed from Bevis et al., 2001). The dark curved shaded area represents the locked segment of the plate boundary between 20 and 50 km depth. The two triangles were used to calculate the infi nitesimal principal strain axes shown below. The fact that both axes are negative means that both are shortening axes. (B) GPS vectors from northern Chile (data from Bevis et al., 2001). The principal strain axes for the triangle shown, which spans Gephart’s (1994) symmetry plane, are also both negative. Four letter codes identify individual GPS stations.



the convergence rate between Nazca and South America has systematically decreased by a factor of 2 since ca. 15 Ma. The magnitude of bound-ary-parallel shortening induced by pre existing curvature is closely related to the convergence rate. Thus, it is not surprising that the major growth of the EW structure is pre–mid-Pliocene.

Trench-Parallel Shortening

Previous investigations have documented Plio-Quaternary trench-parallel shortening farther south in the Andean forearc at 33–37°S (Lavenu and Cembrano, 1999). McCaffrey (McCaffrey, 1994, 1996) has shown that the kinematics of the forearc of subduction zones can be predicted from the relation between the obliquity of plate convergence (relative to the plate boundary) and the obliquity of interplate earthquake slip vec-tors. When looked at in this way, most forearcs experience arc-parallel extension. Northern Chile, however, is one of only two forearcs surveyed by McCaffrey that displays arc parallel shortening, although the large magnitude of the errors in McCaffrey’s analyses does not make this a very robust conclusion. This kinematic analysis is complementary to the elastic mod-eling of Bevis et al. (2001) of a locked, curved plate boundary. Indeed, the arc-parallel shorten-ing predicted from the kinematic analysis arises from the fact that the predicted velocity parallel to the arc not only diminishes linearly from south to north but actually changes sign, just as the arc-parallel components of the velocity vectors in the Bevis model would decrease linearly and change sign. Perfect accord is not achieved, however, as the change in sign in McCaffrey’s model occurs at 23.5°S latitude whereas the change in sign in the Bevis model would occur at the Gephart symmetry axis at 20.5°S.

Of course, as an observational and kinematic analysis, McCaffrey’s model is not required to change sign anywhere. Nonetheless, McCaffrey predicts arc-parallel shortening for all of northern Chile from 30°S to 18°S latitude. As the veloc-ity changes linearly, within error, the strain rate should be constant over this entire length of the forearc. As we have seen, our EW reverse faults are much more restricted latitudinally than the region of predicted arc-parallel shortening.

CONCLUSIONS

We have shown that small but signifi cant shortening parallel to the Nazca-South America plate boundary between 19° and 21.6°S latitude occurred during the Pliocene and may be con-tinuing at present. The scarps produced by this deformation are some of the most prominent topographic features in this segment of the

Chilean Coastal Cordillera. The topographic scarps were initially consistently incised by and later dammed small, well-defi ned drainages in the Coastal Cordillera that are no longer active. The data from one site, Barranco Alto, suggests that tectonic damming of one such drainage produced an internally drained evaporite basin at around 5.6 Ma. This implies, as others have stated as well, that the Coastal Cordillera was wetter at the start of the Pliocene than it is today. Relations around the Río Loa imply either sig-nifi cant uplift of the Coastal Cordillera and/or signifi cant retreat of the Coastal escarpment since the Pliocene. These two processes are probably related to tectonic underplating and subduction erosion, respectively.

The restriction of the EW scarps to the Coastal Cordillera suggests that plate coupling plays an important role in their formation. The latitudinal distribution of the scarps symmetri-cally about Gephart’s symmetry plane implies that processes associated with oroclinal bending are also involved. Though not fi tting our data perfectly, both kinematic analysis of interplate earthquakes and oblique convergence and elas-tic models of locked curved plate boundaries indicate that the northern Chile forearc should experience arc-parallel shortening.

ACKNOWLEDGMENTS

Our ideas about the Coastal Cordillera of northern Chile have been shaped over a number of years by our colleagues, including José Cembrano, Teresa Jordan, Mike Bevis, Constantino Mpodozis, and Jack Love-less. We are indebted to Terry Spell and Kathleen Zanetti of the Nevada Isotope Geochronology Labora-tory for the geochronological analyses reported here. We are grateful to reviewers George Hilley and Adrian Hartley, as well as editors Michael Edwards and Peter Copeland for numerous suggestions that improved the manuscript. The geological work was supported by the U.S. National Science Foundation under grant EAR 0087431 (to Allmendinger and Isacks) and Fundación Andes-Chile grant C-13755-12- (to Gabriel González). The InSAR topographic data set presented here was produced with support from NASA grants NAG5-11424, NAG5-30126, and NSF grant EAR-9706427 (to Isacks).

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