The link between forearc tectonics and Pliocene–Quaternary deformation of the Coastal Cordillera, northern Chile

Journal of South American Earth Sciences 16 (2003) 321–342 www.elsevier.com/locate/jsames

The link between forearc tectonics and Pliocene –Quaternary deformation of the Coastal Cordillera, northern Chile Gabriel Gonza´lez*, Jose´ Cembrano, Daniel Carrizo, Alejandro Macci, Heinz Schneider Depto. de Ciencias Geolo´gicas, Universidad Cato´lica del Norte, Casilla, 1280 Antofagasta, Chile Received 1 February 2003; accepted 1 May 2003

Abstract Pliocene – Quaternary tectonic extension has controlled the structural evolution of the outer forearc of northern Chile. Extensional deformation is documented by large-scale, mostly N– S-striking normal faults and open cracks that run along the forearc. The faults dip to the east and define half-graben geometries across the Coastal Cordillera (208450 – 238300 S). Progressive trenchward tilting of the hangingwalls has produced internal unconformities within the graben and half-graben fill. This suggests that extension has been episodic, at least since the Pliocene. Northwest-striking dextral strike –slip faults occur in the northern part of the extensional fault zone of the Coastal Cordillera (208400 S – 218100 S), where they link N– S normal faults. Hectometric open cracks occur throughout the Coastal Cordillera, some of which are spatially associated with the normal faults. Normal and dextral faults appear to accommodate long-term extensional strain, whereas open cracks are interpreted as tensile fractures formed by either concomitant normal faulting or coseismic extension. The long-term extensional regime may result from buckle folding of the margin, which in turn produces progressive uplift of the coastal area by oceanic plate subduction, followed by continental margin collapse toward the trench. q 2003 Elsevier Ltd. All rights reserved. Keywords: Central Andes; Chile; Forearc; Extensional tectonics; Strike– slip faulting

1. Introduction Continental forearcs are particularly sensitive to deformation induced by plate convergence because their bases (20 – 50 km depth) are seismically coupled with the subducting plates (Ruff, 1996; Ruff and Tichelaar, 1996). The focal mechanisms of earthquakes that occur at such a seismically coupled interface indicate underthrusting caused by the frictional displacement of the subducting lithosphere beneath the overriding plate (Kanamori, 1977; Pacheco et al., 1993). Ruff and Tichelaar (1996) found that the maximum depth of seismic coupling is located just below the coastline of many forearcs of the Circum-Pacific region. This spatial association implies that thrusting along the plate boundary is causally connected to coastal uplift, perhaps due to crustal shortening and thickening driven by plate convergence. Therefore, near-surface deformation should be characterized by orogen-parallel contractional * Corresponding author. Tel.: þ 56-55355952; fax: þ 56-55355977. E-mail address: [email protected] (G. Gonza´lez). 0895-9811/$ - see front matter q 2003 Elsevier Ltd. All rights reserved. doi:10.1016/S0895-9811(03)00100-7

structures that affect the forearc, especially along a strongly coupled plate boundary such as the Chile– Peru´ margin. The continental margin of northern Chile, where the Nazca Plate subducts beneath the South American Plate, is an excellent place to study the interaction between near-surface forearc deformation and deep-seated tectonics driven by plate convergence. In northern Chile, well-exposed structures formed during the Neogene – Recent, provide important information on the longand short-term tectonic processes that have dominated the margin. Although several contributions regarding the local Neogene tectonics of the Antofagasta area and the Mejillones Peninsula have been published (Armijo and Thiele, 1990; Niemeyer et al., 1996; Delouis et al., 1998), systematic geometric and kinematic data for larger portions of the Coastal Cordillera are lacking, and a regional overview of the neotectonics of the northern Chilean Andean forearc has yet to emerge. In this paper, we present new structural field data and provide an integrated model for Pliocene – Quaternary

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deformation and tectonics of the Coastal Cordillera (208450 –238300 S, 500 km segment, Fig. 1) that significantly improves the current understanding of long-term, large-scale processes operating along Andean-type margins.

2. Morphology and uplift of the Coastal Cordillera The Coastal Cordillera is a 1 –2 km high, 25 –45 km wide prominent forearc morphological feature in northern Chile that runs parallel to the trench. The most remarkable feature

Fig. 1. Tectonic map of the Coastal Cordillera between Iquique and Antofagasta. The regional geometry of the Atacama Fault System is shown in the upper left corner (modified from Arabasz, 1971). Trench orientation is indicated in front of the Coastal Cordillera in the Salar Grande area and Mejillones Peninsula. The inset indicates the location of the map shown.

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Fig. 2. View looking to the east along the Coastal Cliff near Salar Grande Area showing the flat erosional surface on top of the Coastal Cordillera. Note the trace of the chomache fault as shown with arrows. In the foreground of the photograph the white color horizon is a pleistocene marine deposit that covers the coastal plain exposed at the foot of the cliff.

of the Cordillera is a very steep, nearly 1000– 2000 m coastal cliff (Fig. 2) that plunges 458 into the sea (Mortimer et al., 1974). The escarpment is largely inactive, but Pleistocene marine terraces exposed at the foothills attest to a formerly intense marine degradation of the coastal cliff profile (Martinez and Niemeyer, 1982; Leonard and Wehmiller, 1991; Ortlieb et al., 1996a). Armijo and Thiele (1990) argue that the coastal cliff is an expression of a major, west-dipping normal fault. However, subsequent work has shown that the coastal cliff is the result of marine erosion acting on an actively uplifting coastline since the Pliocene (Hartley and Jolley, 1995; Niemeyer et al., 1996). The eastern side of the Coastal Cordillera is a transitional zone where a smooth topography marks the gradual boundary of the Cordillera with the Central Depression, a basin filled in some places with more than 1000 m of Tertiary – Quaternary, alluvial fan, fluvial, and lacustrine deposits (Hartley et al., 2000). The top of the Coastal Cordillera is an ancient, wellpreserved erosional surface and testifies to the extremely arid conditions that prevailed in the Atacama Desert since the Early Pliocene (Hartley and Chong, 2002). Such a surface is believed to result from intense erosion of the northern Chilean margin during two time periods. One period relates to the development of the Oligocene – Miocene coastal Tarapaca pediplain (Mortimer et al., 1974). The remnants of the pediplain underlie the sedimentary infill of the Central Depression. According to

K – Ar age determinations from interbedded ignimbrites in the sedimentary fill that range from 21.2 ^ 0.3 to 16.9 ^ 0.3 Ma (Mortimer et al., 1974), they were deposited after the coastal Tarapaca pediplain development. Much younger ages are related to the formation of a widely distributed erosion glacis, which mostly shapes the presentday topography of the Coastal Cordillera, characterized by smooth mountain tops and flat intermountain basins (Fig. 2). K – Ar ages between 2.9 and 6.1 Ma (Table 1) obtained in tuff layers interbedded with alluvial facies of these intermontane basins are believed to approximate the time of formation of this erosion surface. The formation of the modern topography of the Coastal Cordillera ended during the Late Miocene – Early Pliocene. Because the faults described herein cut and displace this surface, they almost certainly represent the youngest deformation phase of the Coastal Cordillera. The coastal cliff in the study area is flanked by a 2 –4 km wide coastal plain of seaward-steeping marine terraces. Well-preserved gastropods and mollusk shells in sediments that cover the marine terraces are Pleistocene in age (Radtke, 1989; Ortlieb et al., 1996a). In many localities, the terraces form a staircase pattern that stretches from near the modern sea level to , 170 m above present-day sea level (asl). South of Iquique, the terraces form five levels, from a maximum altitude of 100 m to a minimum altitude of , 3– 4 m. From Caleta Caramucho to the Rio Loa (Fig. 1), these marine terraces are discontinuously preserved, and

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Table 1 K –Ar age determinations of biotites from volcanic ashes intercalated in alluvial fan deposits of the Salar de Carmen area Sample

Location

SG-007 SC-1 HADM-210 HAGG-151

218030 5100 S, 238310 0000 S, 238430 4800 S, 238430 4800 S,

708070 4200 W 708150 0000 W 708180 4200 W 708200 5800 W

Dated material

%K

%atm. Ar

%Ar nl/g rad (STP)

Age Ma

Reference

Biotite from volcanic ashes Biotite from volcanic ashes Biotite from volcanic ashes Biotite from volcanic ashes

7.918 7.462 7.065 7.121

59 56 78 89

1.715 0.855 1.427 0.801

6.1 ^ 0.3 3.0 ^ 0.2 5.2 ^ 0.5 2.9 ^ 0.5

This work Naranjo (1987) This work This work

their altitude is not . 70 m asl. From Rio Loa to Caleta Michilla, the terraces are blanketed by conspicuous Quaternary alluvial fan deposits. Along this coastal tract, terraces lie at 17 m asl, below the alluvial fans. From Caleta Michilla to Antofagasta, the coastal plain is very well developed and comprises three or four seaward-dipping terraces that reach a maximum altitude of 170 m asl and a minimum altitude of 6 –7 m asl. Just north of Antofagasta, the coastal plain broadens to the west and reaches a maximum width of 12 km to form a morphological landmark of the northern Chilean coastline: the Mejillones Peninsula. The Peninsula has been interpreted as part of the emerged continental shelf (Niemeyer et al., 1996). The marine terraces of the Peninsula are preserved at the top of the highest hills with altitudes in excess of 300 –400 m asl. To determine whether the present-day position of a marine terrace is the result of glacioeustatic sea level fluctuation or tectonic uplift, it is necessary to know the terrace age and the sea level at the time of terrace formation. Worldwide sea level variations during the past 140 ka are better constrained than those for the Middle – Early Pleistocene (Chappel et al., 1996; Pillans et al., 1998). The Late Pleistocene – Holocene sea level fluctuation history is characterized by two maximum interglacial highstands: (1) 5 ^ 2 m asl at 122 ^ 4 Ma (Chappel et al., 1996; Pillans et al., 1998) and (2) , 3 m asl of middle Holocene age (Leonard and Wehmiller, 1991). South of Iquique, a marine terrace roughly coeval with the 122 ^ 4 Ma highstand is preserved at 12 m asl (Radtke, 1989), indicating an average uplift rate of 0.05 mm/a. In Caleta Michilla, Leonard and Wehmiller (1991) document a marine terrace at an altitude of 39– 40 m asl whose origin may be related to the maximum sea level of 122 ^ 4 Ma or to a younger sea level highstand as recent as 80 ka. Subsequent work has demonstrated that the paleosea level at 80 ka was 20 m lower than the present sea level (Chappel et al., 1996). Therefore, this terrace probably corresponds to the 122 ^ 4 Ma highstand, and an uplift rate of 0.3 mm/a can be calculated for the past 122 ka. South of Caleta Michilla, in Hornitos, Ortlieb et al. (1996a) recognize the 122 ^ 4 Ma terrace at an altitude of 30– 36 m asl and calculate an uplift rate of 0.24 mm/a. Uplift rates calculated by Delouis et al. (1998) for the Quaternary near Mejillones Peninsula and Antofagasta range from 0.05 to 0.5 mm/a. The uplift rate data show that the tectonic uplift of the coastal plain is not the same throughout the study area. The amount of Late Pleistocene uplift is much greater in

the Mejillones Peninsula and surrounding areas than north of the Rio Loa.

3. Structural setting of the Coastal Cordillera The Atacama Fault System (AFS), the most important structure of the central Andean forearc (Arabasz, 1971), extends for more than 1000 km between Iquique and La Serena (Fig. 1). Its large-scale geometry was formed during the waning stages of a Late Jurassic –Early Cretaceous magmatic arc that dominated the Coastal Cordillera area (Scheuber and Andriessen, 1990; Scheuber and Gonza´lez, 1999). The activity on the AFS started during the Mesozoic, when oblique subduction of the Farallon Plate beneath the South American Plate took place (e.g. Scheuber and Andriessen, 1990; Scheuber and Gonza´lez, 1999). Large brittle structures of the AFS (more than 60 km in length) were formed by sinistral strike – slip (Scheuber and Andriessen, 1990). The inherited plan view geometry of the AFS consists of a series of strike – slip duplexes formed by NS-striking faults and NW-striking splay faults (Fig. 1). The northern edge of the AFS shows a horsetail architecture, where three NW faults that cut the Coastal Cordillera splay off the master faults of the AFS (Fig. 1). Cenozoic reactivation of the AFS structure is suggested by the prominent scarps (30 –100 m in height) that control the horst and graben topography of the Coastal Cordillera (Arabasz, 1971; Okada, 1971; Herve´, 1987; Naranjo, 1987). This fault-controlled morphology has been regarded as evidence in support of Cenozoic reactivation of the AFS in an extensional regime (Herve´, 1987).

4. Seismological framework and coseismic deformation The seismicity of the Coastal Cordillera is strongly dominated by subduction earthquakes that occur along the Wadati – Benioff Zone (Delouis et al., 1996). Focal mechanisms document low-angle thrusting that results from the frictional displacement of the Nazca Plate beneath the South American Plate (Delouis et al., 1996; Ruegg et al., 1996). Published seismic data for the area between Iquique and Antofagasta show that the coupling zone starts at a depth of nearly 20 km and extends to 50 km below the Coastal Cordillera (Delouis et al., 1996; Comte et al., 1999). The shallow locked part of the Wadati – Benioff Zone has

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a dip angle of nearly 208. The last great earthquake in this area, on July 30, 1995, was the Mw ¼ 8.1 Antofagasta event (Ruegg et al., 1996) located beneath the Coastal Cordillera at 23 –258 lat. S. Focal mechanisms show that the slip vector was essentially parallel to the convergence direction, which indicates that no slip partitioning occurred (Ruegg et al., 1996). In turn, this suggests elastic behavior by the forearc crust and agrees with McCaffrey (1994), who proposes an elastic rheology for the north Chilean forearc. Coseismic, ENE –WSW, near-surface crustal movement was revealed by geodetic data following the Antofagasta earthquake (Ruegg et al., 1996; Klotz et al., 1999). The displacement magnitude increased westward, reaching 90 cm in the Coastal Cordillera and 10 cm inland, 300 km east of the Coastal Cordillera (Klotz et al., 1999). Transforming the horizontal displacement gradient into strain, it is possible to demonstrate that the forearc suffered extension in a broad area covering the Coastal Cordillera and regions 300 km inland. The origin of this short-term extensional strain is related to the elastic rebound of the overriding plate during and after subduction earthquakes. Geodetic data show that elastic shortening of the overriding plate (N71E-directed) along the central Andean forearc accumulates during the interseismic period between large subduction earthquakes (Bevis et al., 1999; Klotz et al., 1999), whereas elastic rebound occurs during the coseismic period. The shortening direction is roughly parallel to the present-day convergence vector; the elastic rebound parallels the Nazca-South American Plate convergence direction of N78 – 79E (DeMets et al., 1994) but in an opposite sense.

5. Neogene –Quaternary deformation in the Coastal Cordillera Neogene –Quaternary near-surface faulting along parts of the Coastal Cordillera has already been described (e.g. Armijo and Thiele, 1990; Delouis et al., 1998; Reijs and McClay, 1998). The most spectacular expression of this young upper crustal deformation is observed in three areas of the Coastal Cordillera (Fig. 1): (1) the Salar Grande area, 40 km south of Iquique; (2) the Mejillones Peninsula, 20 km northwest of Antofagasta; and (3) the Salar del Carmen area, directly east of Antofagasta. In these three regions, we conducted detailed structural observations to assess the nature, geometry and kinematics of near-surface crustal deformation. We present our observations related to faultinduced morphology and kinematics and provide arguments to constrain the age of deformation. 5.1. The Salar Grande area 5.1.1. Fault geometry and kinematics The large-scale fault arrangement in the Salar Grande area consists of two NW-striking strike – slip faults and

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several NS- to NNE-striking normal faults (Fig. 3). The most important NW-striking faults are the Salar Grande Fault and the Chomache Fault; both structures splay toward the west from NS-trending faults (Fig. 3). Other faults, exposed along the southeastern border of the Salar Grande, are E – W reverse faults. The kinematics of these faults has been documented recently, and their complete description forms part of another contribution (Allmendinger et al., in prep.). The Chomache Fault has an , 18 km long, straight trace along the western margin of the Coastal Cordillera. At its northern end, the Fault locally cuts the coastal cliff to form a concave fault dipping 758 to the east. At its southern end, the fault surface is not exposed, but the straight fault traces suggest a vertical dip. The fault trace is marked by a surface rupture across Pliocene – Pleistocene alluvial fans; the lateral offset of alluvial fans and ravines indicates 3– 25 m of dextral slip (Fig. 4a). Pull-apart structures (, 1 km long £ 100 m wide) occur along releasing bends, whereas push-up swells occur along restraining bends; both are compatible with dextral displacement. Dextral strike –slip movement along the fault is also supported by subhorizontal striae and mesoscopic kinematic indicators (Fig. 3, sites b, c). The Salar Grande Fault runs subparallel to the Chomache Fault (Fig. 3) along the central part of the Coastal Cordillera and cross-cuts the Salar Grande Basin, which is filled with 100 – 150 m pure halite (Chong et al., 1999). The Fault has a sharp linear trace, with subsidiary structures that locally form fault-parallel elongate lenses. A 10 – 50 m high, southwest-facing scarp is exposed in the central part of the Salar Grande where it cross-cuts the subhorizontal surface of the halite infill. The fault trace forms a wide curved trace concave to the east, suggesting that the Fault dips to the east at a high angle. Close to the Cerro Rojo mine (Fig. 3, site a), the scarp is not developed, and the Fault is subvertical. In this part, several ravines flowing from the Carrasco Range to the piedmont form conspicuous valleys that show right-lateral offsets as they approach the Fault. The horizontal separation magnitude ranges from 20 – 200 m (Fig. 5). A dextral kinematic indicator is documented by a drag fold exposed along the northern part of the Salar Grande Fault, which affects Late Jurassic – Early Cretaceous red beds (Fig. 5). The fault gouge has subhorizontal striae (Fig. 3, site a), and dextral strike – slip kinematic indicators, such as S –C fabrics and ellipsoidal clasts with major axes oriented parallel to the S surface of the gouge, are exposed near the Cerro Rojo mine. These kinematic indicators are consistent with the displacement sense suggested by morphological analysis. In the central part of the Salar Grande, the easterly concave fault trace and southwest-facing scarp imply that the net slip along this fault tract has both dextral and reverse components (Fig. 3). The 8 –36 km long, NS-striking faults of the Salar Grande area exert control over the morphology of the Coastal Cordillera by defining an asymmetric half-graben

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Fig. 3. Topographic model of the Salar Grande area displaying the main faults. Fault-slip data presented as fault plane solutions are shown. The white and black areas correspond to contractional and extensional quadrants, respectively. Rose diagrams show open crack orientations.

morphology. The half-grabens are located east of the faults and flanked by 50 –100 m east-facing scarps (Fig. 3). The Coastal Cordillera pediplain is preserved on top of the grabens and tilted 10 – 158 to the west. This arrangement

resembles a large-scale domino geometry and thus indicates that these structures are normal faults. In most cases, fault surfaces are not well exposed, so they do not provide any direct evidence of their kinematic

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Fig. 4. Dextral kinematics of the Chomache Fault derived from morphological analysis. (a) Right lateral offset of ravines across the Chomache Fault; note the shutter ridge at the foot of the mountain front; view to the northeast. (b) Profile view of the Punta de Lobos Fault forming the contact between Late Miocene– Early Pliocene alluvial deposits (right side) with Mesozoic diorites (left side); view to the north.

character. The Punta de Lobos Fault, the best exposed NSstriking fault in the region, dips 828 east and constitutes the boundary between a 1000 m high, NS-trending mountain range and an adjacent alluvial basin to the east (Fig. 4b). The Fault has several parallel fault scarps that cut into the surface of the alluvial infill. Its exposures show down-dip striae (Fig. 3, sites d –f) and have outcrop-scale kinematic indicators of normal displacement. A spectacular, complex system of open cracks affecting both the sedimentary cover and the bedrocks occurs in the Salar Grande area (Fig. 6). The cracks are hectometric in length, 0.3 –5 m in width, and 1 –7 m in depth. They strike mainly NS, NNW, and, less commonly, NE and

EW (Fig. 3, sites g– l). The NS open cracks occur locally on the footwall of NS-striking normal faults, such as the Punta de Lobos and the Geoglifo Faults (Fig. 6). The NNW cracks are developed in clusters in low topographic zones that affect the alluvial graben infill flanked by the NS-striking normal faults. The NE- and EW-striking cracks occur in clusters mainly in the highest areas around the Salar Grande. Some cracks were formed in the uplifted part of old EW-oriented fault scarps, whereas others are not spatially related to faults. By affecting both highlands and lowlands, open cracks account for an alternative origin that is related to tectonic processes or topographically controlled gravitational collapse.

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Fig. 5. Aerial photograph of the northernmost part of the Salar Grande Fault showing a drag fold of Upper Jurassic– Lower Cretaceous red beds. The arrows indicate sites with dextral offsets of drainage channels. See Fig. 3 for location.

5.1.2. Timing of deformation Detailed field observations suggest that the faults of the Salar Grande area have different relative ages. For example, some faults form the morphological boundary between ranges and basins but show no fresh ruptures on the surface. Other faults, in contrast, are morphologically expressed as fresh ruptures that clearly break the piedmont of the fault-controlled ranges; these faults are obviously the younger of the two types. In addition to fault morphology, we use some regional-scale morphological units to constrain the relative age of the structures, including the Late Miocene – Early Pliocene surface of the Coastal Cordillera and the Plio-Pleistocene coastal cliff. We also use K –Ar isotope age determinations of biotite from tuff lenses interbedded with the faulted alluvial fan deposits. Because both the Chomache Fault and the Salar Grande Fault show well-defined fault traces, displaced channels, and well-preserved fault scarps, it is likely that the dextral displacements along these faults occurred in Pliocene or Quaternary times. The Chomache Fault cuts the youngest alluvial fan deposits and displaces the Plio-Pleistocene coastal escarpment (Niemeyer et al., 1996), which suggests that dextral slip along this fault is of Pleistocene or Holocene age. The Salar Grande Fault has a more strongly eroded morphological scarp and does not show a fresh rupture, which indicates that dextral displacement along this fault is older than that of the Chomache Fault.

Fig. 6. General view of the Geoglifo Fault looking south. The fault scarp and open crack cluster, localized in the uplifted fault block, are shown. Center-right, another fracture cluster is related to a different fault.

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The NS-striking faults exhibit different degrees of degradation that indicate different ages of activity. For example, those located in the northernmost part of the Salar Grande are more eroded and do not show fresh rupture surfaces across alluvial fan deposits, and the cracks are filled with eolian and alluvial sediments. In contrast, those faults located near the coastal cliff have mildly eroded fault scarps and unfilled cracks, both of which attest to a younger age for the normal slip. The Punta de Lobos Fault deforms alluvial fan deposits with interbedded tuffs. The oldest tuff layer yields a 6.1 ^ 0.3 Ma K – Ar age (Table 1), which indicates a post-Late Miocene age. 5.2. Mejillones Peninsula 5.2.1. Fault geometry and kinematics In the Mejillones Peninsula, Miocene – Pleistocene marine deposits fill two half-grabens flanked by the Caleta Herradura and Mejillones Faults (Fig. 7). The footwalls west of the structures form two conspicuous horsts, the Jorgino and Mejillones horsts (Fig. 7), that consist of Paleozoic metamorphic rocks and Paleozoic – Jurassic plutonic rocks. Pliocene – Pleistocene marine deposits are preserved on top of these horsts (Fig. 7) and Miocene – Pleistocene marine deposits in the infill of the grabens (Ibaraki, 1990; Koizumi, 1990; Niemeyer et al., 1996). The age of the sedimentary deposits has been constrained by micro- and macrofauna (Ibaraki, 1990; Koizumi, 1990; Niemeyer et al., 1996). The offset of the stratigraphic markers shows that eastern blocks are downthrown relative to western blocks throughout the area. The Caleta Herradura Fault is exposed along the eastern boundary of the Jorgino horst. At the surface, the Fault dips 558 to the east; however, 5–308 westward tilting of the Miocene–Pliocene graben infill east of the Fault suggests that the fault dip decreases with depth (Fig. 8). Tilting of the halfgraben infill can result from down-dip displacement of the hangingwall on an east-dipping fault with listric geometry. The sedimentary strata in the half-graben are thick near the Caleta Herradura Fault and thinner away from the Fault, in support of the listric geometry. This fault geometry is confirmed in that only the marine deposits in the hangingwall experienced tilting, whereas the sedimentary deposits and marine terraces in the footwall maintain their subhorizontal bedding. Work in progress also shows that the Caleta Herradura Fault has a subhorizontal detachment 2–3 km below the half-graben surface (Gonza´lez et al., in prep.). It is possible to estimate the magnitude of vertical displacement using Pliocene and Pleistocene marine deposits preserved at the tops of the Jorgino horst and graben infill. Along the southern part of the Mejillones Peninsula, the Caleta Herradura Fault displaces Pleistocene marine deposits (Fig. 7, site m). Vertical separation of these deposits produces a cumulative vertical offset of nearly 100 m. In the central part (Fig. 7, site n), the vertical offset, according to Pliocene and Pleistocene marine deposits, is

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, 150 m; in the northern part (Fig. 7, site o), the vertical offsets of the Pleistocene deposits reach nearly 150 m. By using the unconformity that separates the Paleozoic basement from the early Miocene, we find an approximately 400 m vertical separation. The nearly down-dip fault striae (758 rake) indicate that the Caleta Herradura Fault has a normal slip with a very small left-lateral component. Thus, the documented vertical offsets are close to net displacements. The Mejillones Fault has the same extensional character as the Caleta Herradura Fault; slip magnitude across the fault is marked by the offset of Pliocene –Pleistocene marine deposits. In the southern part (Fig. 7, site p), the Fault displaces Pliocene strata with a vertical separation of , 70 m. In the middle segment (Fig. 7, site q), the vertical separation of Pleistocene sediments is nearly 250 m, whereas in the northern part (Fig. 7, site r), the separation of Pleistocene deposits is , 400 m. The northward increase in the vertical separation indicates that the southern part of the fault is closer to the tip point. The sediments that fill the Mejillones Graben contain several decametric cracks filled with sand exposed along a 5 –7 m high cliff near the present-day coastline. Some of these cracks are hybrid normal faults. On the surface of the Mejillones Graben, several beachline ridges, parallel to the present-day coastline, are located at 10 –220 m asl (Fig. 6). At , 160 m asl, Ortlieb et al. (1996a) find warm-water paleofauna assigned to 400 Ky. In the southern part of Mejillones Peninsula, Ortlieb et al. (1996b), using the same paleontological fauna, assign the oldest beach ridges (50 m) to isotope stage 11 (400 Ky). For the southernmost terrace, Ortlieb (1995) dates mollusk shells and obtains a 300 Ky age (U –Th series). The concave-to-the-sea, concentric geometry of the beach ridge traces in both parts of the Peninsula indicates that, during the Pleistocene, the sea abandoned the land away from the central part of the Peninsula in both the north and south. This coincides well with an uplift of the central part of the Mejillones Peninsula. That the Mejillones Fault increases its displacement from the central part of the Peninsula toward the north suggests that normal faulting, sea regression, and uplift of the Mejillones Peninsula were interlinked processes. Considering that the beach ridges are better developed on the top of the footwall, the Mejillones Peninsula and its associated normal faults must have experienced uplift driven by some deep-seated process. The preservation of the oldest beach ridges at an altitude of 220 m asl suggests that the accumulated uplift of the surface of the Mejillones Graben is close to 200 m (Ortlieb et al., 1996b). 5.2.2. Timing of deformation Internal unconformities in the graben infill of the Caleta Herradura Fault show progressive westward tilting of the marine deposits. The tilting amount decreases from 308 in the Miocene strata to 58 in the Pliocene – Pleistocene

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Fig. 7. Digital elevation model of the Mejillones Peninsula including the most conspicuous faults. Fault-slip data of the Caleta Herradura Fault presented on the left side as fault plane solutions, and fault-slip data of the Mejillones Faults shown on the right side. Age estimated on the basis of paleontological data by Ortlieb et al. (1996b).

deposits and suggests that the infilling of the half-graben basins was syntectonic. Younger alluvial fans derived from the flanks of the Jorgino and Mejillones horsts cover the Pleistocene marine sediments and are cut by the Caleta

Herradura and Mejillones Faults, which indicate that the structures were active during Late Pleistocene times. Therefore, normal faulting was active from the Pliocene to the Late Pleistocene.

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Fig. 8. Panoramic view of the Caleta Herradura Fault to the southwest. Miocene– Pleistocene marine sedimentary graben infill shown in the center and left. On the right side, flat surfaces at the top of Jorgino horst represent marine abrasion platforms of Pliocene age. An internal unconformity, on the right, separates Miocene diatomites from Pliocene sandstones.

5.3. Salar del Carmen 5.3.1. Fault geometry and kinematics A sharp tectonic lineament, the Salar del Carmen Fault, occurs on the eastern side of the Coastal Cordillera between Taltal and Salar Grande (Arabasz, 1971; Naranjo, 1987). Detailed field observations in the Salar del Carmen area show that this Fault is actually formed by two nearly parallel strands that affect the eastern piedmont of the Coastal Cordillera (Fig. 9). The western strand is the Sierra del Ancla Fault, and the eastern strand is the Salar del Carmen Fault. The Sierra del Ancla Fault marks the morphological boundary between the mountain front of the Coastal Cordillera and its eastern piedmont, which is formed by alluvial fan deposits (Fig. 9). The fault strikes , N20E, dips 72– 848E, and extends for more than 23 km from La Negra to 10 km north of Salar del Carmen (Fig. 9). Detailed field observations show that alluvial fan deposits are cut and displaced down-dip to form a 0.3– 1.5 m high, east-facing scarp. Well-preserved down-dip striae are consistent with the normal nature of the Sierra del Ancla Fault. The Salar del Carmen Fault affects alluvial fans of the eastern piedmont of the Coastal Cordillera (Fig. 10). Displacement along the fault has formed a 0.5 – 9 m high, east-facing scarp. The Salar del Carmen Fault consists of three segments (Fig. 9) joined by relay ramp structures. The southern segment, located between La Negra and the Salar del Carmen Basin, strikes N –S to N10E. Several small scarps develop along this segment, and their displacement is synthetic to the main rupture. The central,

shortest segment strikes N15E and forms the western boundary of the Salar del Carmen Basin. The northern segment, 2 km north of Salar del Carmen, strikes N15E and extends for more than 30 km. Along these three segments, the vertical separation of alluvial fans indicates that fault kinematics is dominated by east-down slip. Only the northernmost part of the fault scarp faces the west, so the slip is west-down. Stratigraphic markers, such as volcanic tuff layers in the oldest alluvial fan deposits and bedding of these deposits, show that displacement is mainly normal (Fig. 10a). Cross-sections of the Sierra del Ancla Fault and the southern segment of the Salar del Carmen Fault show that the alluvial fan deposits and present-day surface are tilted 10 – 158 to the west (Fig. 10b). This tilting can be interpreted as roll-over structures related to normal faulting along faults whose dip progressively shallows with depth. Small-scale reverse faults striking N15E and dipping 65 – 758 west (with 0.5 – 1 m vertical offset) occur locally in the hangingwall of the Salar del Carmen Fault. These reverse faults merge from a blind, subvertical fault parallel to the Salar del Carmen Fault. The alluvial fan strata above the faults form an open fold, which suggests that these structures are kinematically coupled with the upward propagation of the subvertical fault tip line, similar to the extensional forced fold described by Withjack et al. (1990). Armijo and Thiele (1990) report a 60 –100 m sinistral offset along the northern segment of the Salar del Carmen Fault. They argue that, for each of three consecutive ravines crossing the fault, the younger alluvial fans were located systematically to the south with respect to the older alluvial

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Fig. 9. Digital elevation model of the Salar del Carmen area, eastern side of Coastal Cordillera. Main faults are the Sierra del Ancla and Salar del Carmen Faults. Fault-slip data presented as fault plane solutions are shown on the left.

fans on the eastern fault block. Delouis et al. (1998) explain Armijo and Thiele’s (1990) observation as a topographic effect. Our field observations confirm Delouis et al.’s (1998) arguments. Along the Salar del Carmen, there are several 10 –100 m long, 2 m deep, open subvertical cracks oriented subparallel

to the main faults. The cracks are located at the relay ramp between the segments of the Salar Carmen Faults and at its southern termination. Other cracks, filled with clastic debris, occur in subvertical sections of the alluvial fan deposits, which suggests that they were episodically formed during alluvial fan construction.

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Fig. 10. Normal to subvertical separation marked by strata of alluvial fan deposits in the Salar de Carmen Fault. (a) Quarry view to the north showing alluvial fan deposits cut by the fault; white layer is a tuff. Note the westward tilting of the hangingwall on the right side of the picture. (b) Looking to the northwest, a fault scarp showing the vertical separation of alluvial fan deposits. The hangingwall exhibits a roll-over structure marked by bending of the alluvial fan deposits.

5.3.2. Timing of deformation In the Salar del Carmen area, faulting was probably active during the construction of the eastern piedmont, as shown by the oldest inactive alluvial fans, which have their

apices in the mountain front, whereas the youngest alluvial fans have their apices localized along the main scarp. Furthermore, the Salar del Carmen Fault probably was active during the construction of the youngest alluvial fan,

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as indicated by a faulted debris slope from the main scarp. Detailed mapping of trenches along and across the Fault reveals that deformation was episodic. Colluvial wedges derived from the degradation of the main scarp are affected by normal faults. The age of the youngest slip event of the Salar del Carmen Fault was constrained using K – Ar determinations of biotite from tuff layers interbedded in the oldest cross-cut alluvial fans. Two ages, 2.9 ^ 0.5 Ma and 5.2 ^ 0.5 Ma (Table 1), were obtained; another K – Ar biotite age of 3.0 ^ 0.2 Ma has been documented by Naranjo (1987). These data indicate that normal faulting took place during the Late Pliocene – Quaternary. In that the cracks close to the Salar del Carmen Fault remain open and the area is affected by regional-scale mudslides due to the unusually heavy rains associated with El Nin˜o events (Vargas et al., 2000), we suspect that the open cracks devoid of sediments are very young (, 10 Ky).

6. Young deformation in the Rio Loa Canyon Between the Salar Grande and Salar del Carmen areas, the AFS shows many scarps that form a horst and graben topography in the central part of the Coastal Cordillera. Near the Rio Loa Canyon the main branch of the AFS has NS-striking fault trace on the surface of Neogene fluvial sediments. East of this fault trace, several NS-striking faults with vertical-normal separation can be found along the east – west canyon of the Rio Loa (Fig. 11). The faults dip 70– 808 to the east and have 2– 20 m down-east vertical offsets. This vertical displacement is compatible with the Coastal Cordillera’s placement west of the Central Depression. Normal faults in the Rio Loa Canyon show that extensional deformation is also present in other areas of the Coastal Cordillera and compatible with a regional extensional strain field younger than Early Pliocene in age.

Fig. 11. To the south of the normal fault exposed in the vertical profile of the Rio Loa Canyon. Arrows show the displaced unconformity that separates the underlying Paleozoic basement from alluvial and lacustrine neogene deposits.

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7. Fault strain compatibility Using the attitude of the fault plane, the orientation of the slip direction, and the sense of slip, we apply the kinematic graphical reconstruction of Marrett and Allmendinger (1991) to obtain the P- and T-axes (Table 2)-the principal incremental shortening axis ðPÞ and the principal incremental extensional axis ðTÞ: The P- and T-axes are located in the movement plane at a 458 angle to the nodal planes of fault plane solutions. In a stereoplot, P- and T-quadrants are defined by the region between the two nodal planes of fault plane solutions. A first-order distribution of the strain axes can determine whether the faults are kinematically compatible. In the Salar Grande area, the Salar Grande Fault (Fig. 3, site a), the Chomache Fault (Fig. 3, sites a, c), and the Punta de Lobos Fault (Fig. 3, sites d, e) show well-developed striae. The T-axes obtained from the two dextral strike – slip faults (Salar Grande and Chomache) have a subhorizontal, EW orientation. The attitude of the P-axes is nearly subhorizontal and trends NS (Fig. 3, sites a– c). The linked Bingham distribution of the tensor axes for the right lateral faults yields a N898E-trending, subhorizontal T-axis and a N01W-trending, subhorizontal P-axis (Fig. 12a, Table 2). For the Punta de Lobos Fault, the P – T analysis shows that the T-axes are subhorizontal and trend EW. In contrast to the Salar Grande Fault, the P-axis derived from the Punta de Lobos Fault is subvertical, as determined by the normal character of the Fault (Fig. 3, sites d, e). In other normal faults, we calculate a similar orientation of the T- and P-axes (Fig. 3, site f). Applying the Bingham analysis to the whole normal fault population, we obtain a N788E-trending, horizontal T-axis (Fig. 12b, Table 2). This attitude is similar to that yielded by the dextral faults. In the Mejillones Peninsula, fault striae occur on the Caleta Herradura Fault (Fig. 7, sites s, t) and the Mejillones Fault. Striae are poorly developed (s4) or not exposed at all on other minor NS-striking faults, but fault separation of stratigraphic markers shows dominantly normal slip (Fig. 7, sites s1 – s3 and t1 – t2). In the Caleta Herradura Graben, these faults define a well-developed conjugate fault system. Because of the poor exposure of the striae on these conjugate faults, we reconstructed the theoretical position of the average slip vector for the entire fault population, as follows: (1) we plotted the minor conjugate fault for which we knew the attitude of the fault plane and fault separation from stratigraphic markers; (2) in a contour diagram, we obtained the average fault orientation for each of the two groups of conjugate minor faults; and (3) because the intersection of these two averaged faults defines a B axis, we determined that the slip vector is located 908 from it. This procedure is valid only if plane strain is conserved. The calculated T-axes for the Herradura and Mejillones Faults are subhorizontal and trend N918E and N98E, respectively (Table 2). The calculated T-axes orientation of faults exposed in the Caleta Herradura Graben correspond

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well to the overall subhorizontal EW extension (Table 2). For the faults exposed in the Mejillones Graben, we obtained a Bingham distribution with shallowly westplunging T-axes. When applying the Bingham statistical analysis to the whole fault population of the Mejillones Peninsula, we obtained a tensor axis distribution characterized by a N94E-trending subhorizontal T-axis and subvertical P-axis (Table 2). Striae occur on the Salar del Carmen Fault (Fig. 9, sites u1, u3), Sierra del Ancla Fault (Fig. 9, site y), and other minor faults (Fig. 9, site u2) of Salar del Carmen. The Bingham analysis for the faults of the Salar del Carmen area shows that deformation is consistent with a N948E-trending, subhorizontal T-axis (Table 2; Fig. 9, sites u1 – u3, y). By using Bingham analysis, we calculated the tensor axes for the Salar Grande fault slip data group (Fig. 12c) and a combined Mejillones Peninsula and Salar del Carmen fault slip data group (Fig. 12d). For the first group, the results show a N858E-trending, subhorizontal T-axis. For the second group, a N94E-trending, subhorizontal T-axis was obtained. The two T-axes populations cannot be distinguished at the 95% confidence level, which suggests that the deformation of the two areas corresponds to the same EW extension event. Using crack orientation, we estimated the T-axes relative to their formation. For each crack, the fracture plane pole corresponds to the orientation of the T-axes. In the Salar Grande area, the fracture orientation defines a maximum pole located between N60 – 70E, with a second maximum oriented N80 – 90E (Fig. 12e). In the Salar del Carmen, the maximum is located between N100 – 110E, with a second maximum in the azimuth N110 –120E (Fig. 12f). Fault slip data in the three selected areas display evidence of kinematic compatibility in a bulk, EW extensional regime. The most relevant question is the role of the NW-striking dextral strike – slip faults of the Salar Grande area in bulk deformation. The following field observations support the idea that strike – slip displacement is kinematically related to EW extension: (1) The Salar Grande and Chomache Faults do not cross-cut the Salar del Carmen Fault at the southern tip of the Salar Grande; (2) the Chomache and Salar Grande Faults connect two NS-striking normal faults with opposite dip sense; and (3) the T-axes obtained for the strike – slip and normal faults have similar attitudes and orientations. These observations strongly suggest that the normal and strike – slip faults form a linked fault system. Alternatively, the NS-oriented P-axes obtained from the NE-striking dextral faults could mean that their orientation is related to a N –S shortening of this part of the Coastal Cordillera near Salar Grande. In this scenario, dextral slip along northwest faults could be kinematically coupled with the observed EW extension accommodated by N –S normal faults. The comparison between the T-axes obtained from the fault slip data and those derived from crack orientation show

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Table 2 Kinematic data from fault populations of the three selected areas of northern Chile Site (coordinates)

Fault

Striae

Sense

P

T

MF MF MF MF

320 78 140 75 130 75 306 65

140 00 311 29 305 19 126 00

RL RL RL RL

184 09 004 32 353 25 168 17 358 08

276 09 268 09 262 03 264 17 267 09

SF SF SF SF SF

138 80 153 72 129 75 123 65 150 75

141 15 153 00 305 15 127 09 320 34

RL RL RL RL RL

185 03 017 13 352 21 348 11 015 35

094 18 109 13 262 00 083 24 276 12

MF

120 75

125 19

RL

168 03 002 15 359 10

077 25 094 08 89 01

SF SF SF SF SF SF SF SF SF SF SF SF SF

332 40 344 55 005 45 004 25 355 40 000 75 356 30 000 50 348 75 352 40 345 55 352 32 357 60

037 37 074 55 088 45 094 25 085 40 056 72 086 30 090 50 032 69 082 40 058 54 082 32 123 55

NL NL NL NL NL NL NR NL NL NR NL NR NR

345 76 254 80 002 88 094 70 085 85 284 59 086 75 270 85 278 57 082 85 289 77 082 77 226 68

228 07 074 10 092 00 274 20 265 05 082 29 266 15 090 05 066 29 262 05 068 10 262 13 101 13

Site e 218030 5100 –708070 4200 Linked Bingham statistic for Punta de Lobos Fault

MF

342 82

020 77

NL

264 52 288 87

063 36 081 03

Others faults Site f 218040 2200 –708060 1800 218040 2200 –708060 1800 Linked Bingham statistic for Punta de Lobos and other faults General Linked Bingham statistic for Salar Grande area

SF SF

020 75 015 86

071 71 018 38

NL NL

306 58 324 33 292 81 341 67

100 29 068 23 078 07 85 06

MF

010 65

073 62

NL

303 68

091 19

SF SF

194 82 034 54

353 69 084 47

NR NL

082 49 001 68 055 64

301 34 106 06 292 14

SF SF

157 78 006 64

322 50 032 42

NR NL

030 44 325 52 004 53

275 24 067 09 261 07

SF

348 70

066 70

NL

265 65

075 25

Salar Grande ðn ¼ 26Þ Salar Grande Fault Site a 208520 2200 S–708040 4800 W 208520 2200 S–708040 4800 W 208520 2200 S–708040 4800 W 208520 2200 S–708040 4800 W Linked Bingham statistic Chomache Fault Site b 218030 5100 –708070 0700 218030 5100 –708070 0700 218030 5100 –708070 0700 218030 5100 –708070 0700 218030 5100 –708070 0700 Site c Linked Bingham statistic Linked Bingham statistic for Salar Grande and Chomache Fault Punta de Lobos Fault Site d 218030 5100 –708070 4200 218030 5100 –708070 4200 218030 5100 –708070 4200 218030 5100 –708070 4200 218030 5100 –708070 4200 218030 5100 –708070 4200 218030 5100 –708070 4200 218030 5100 –708070 4200 218030 5100 –708070 4200 218030 5100 –708070 4200 218030 5100 –708070 4200 218030 5100 –708070 4200 218030 5100 –708070 4200

Mejillones Peninsula ðn ¼ 22Þ Caleta Herradura Fault Site s 238120 3400 –708340 28 Faults in the Herradura Graben Site s1 238120 3000 –708340 2100 n ¼ 22 238120 3000 –708340 2100 n ¼ 37 Linked Bingham statistic Site s2 238120 2400 –708340 1200 n ¼ 23 238120 2400 –708340 2100 n ¼ 24 Linked Bingham statistic Site s3 238120 1000 –708330 5900 n ¼ 14

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Table 2 (continued) Site (coordinates)

Fault

Striae

Sense

P

T

SF

166 62

265 62

NR

066 73 296 85 020 74

259 17 077 04 270 06

Site s4 238270 5200 S – 708310 0200 W

SF

025 40

115 40

NR

206 90

115 85

Mejillones Fault Site t 238060 5100 S – 708290 3200 W

MF

012 55

093 55

NL

301 79

98 10

SF SF SF SF SF SF SF

358 70 025 45 000 70 000 80 010 70 354 85 008 80

041 62 115 45 090 70 090 80 062 65 129 83 098 80

NL NR NL NL NL NR NL

297 60 207 90 270 65 270 55 303 62 259 50 278 55 275 63

073 23 115 00 090 25 090 35 088 23 089 40 098 35 092 27

SF SF SF SF SF SF

010 80 180 82 000 80 180 74 180 70 180 90

030 63 343 64 090 80 270 74 270 70 270 90

NL NR NL NR NR NR

307 49 064 47 270 55 090 61 090 65 090 45 090 55 317 79 309 74

080 31 291 32 090 35 270 29 270 25 270 45 270 35 094 08 94 13

MF MF MF MF

006 80 002 75 020 80 337 88

096 80 073 74 083 79 150 75

NR NL NL NR

276 55 279 60 296 55 232 45 302 66

096 35 088 30 106 35 081 41 094 21

SF SF

185 77 175 75

296 76 299 72

NR NR

088 58 071 59

279 32 273 29

Site u3 238480 4300 S – 708190 4200 W 238480 4300 S – 708190 4200 W 238490 2700 S – 708190 3700 W 238490 2700 S – 708190 3700 WS 238490 2700 S – 708190 3700 W 238490 2700 S – 708190 3700 W Linked Bingham statistic

MF MF MF SF SF MF

005 77 355 75 345 90 160 85 170 85 170 85

074 76 051 72 165 75 219 84 260 85 194 78

NL NL NL NL NR NL

282 58 279 59 240 43 073 50 080 50 092 49 066 69

091 32 077 29 289 43 247 40 260 40 250 39 264 20

Sierra del Ancla F. Site y 238400 4500 S – 708190 0600 W 238400 4500 S – 708190 0500 W 238400 4500 S – 708190 0600 W 238400 4500 S – 708190 0600 W 238400 4500 S – 708190 0600 W Linked Bingham statistc General Linked Bingham statistic for the Salar del Carmen area

MF MF MF MF MF

010 65 005 72 012 88 020 75 020 75

073 62 049 65 158 86 066 70 082 73

NL NL NR NL NL

303 68 300 59 279 47 309 58 301 59 298 59 302 66

091 19 081 25 105 43 099 29 104 30 096 29 0094 21

238120 1000 – 708330 5900 n ¼ 6 Linked Bingham statistic Linked Bingham statistic for the faults in the Graben

Faults in the Mejillones Graben Site t1 23850 5000 S – 708260 4600 W 23850 5000 S – 708260 4600 W 23850 5000 S – 708260 4600 W 23850 5000 S – 708260 4600 W 23850 5000 S – 708260 4600 W 23850 5000 S – 708260 4600 W 23850 5000 S – 708260 4600 W Linked Bingham statistic Site t2 23850 5000 S – 708280 1300 W 23850 5000 S – 708280 1300 W 23850 5000 S – 708280 1300 W 23850 5000 S – 708280 1300 W 23850 5000 S – 708280 1300 W 23850 5000 S – 708280 1300 W Linked Bingham statistic Linked Bingham statistic for Mejillones Peninsula Linked Bingham statistic for Mejillones Peninsula and Salar del Carmen Salar del Carmen ðn ¼ 17Þ Salar del Carmen Fault Northern Segment Site u1 238350 5600 S – 708190 3400 W 238350 5600 S – 708190 3400 W 238350 5600 S – 708190 3400 W 238350 5600 S – 708190 3400 W Linked Bingham statistic Southern segment Site u2 238400 4100 S – 708240 0500 W 238400 4100 S – 708240 0500 W

Fault attitude follows the right-hand rule; for striae, T- and P-axes plunge directions and plunge. MF ¼ main fault and SF ¼ secondary fault.

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Fig. 12. (a –d) Contour plots for the T-axes obtained from fault slip data and Bingham statistics for strain axes. (a) Dextral faults of Salar Grande ðn ¼ 10Þ; (b) normal faults of Salar Grande ðn ¼ 17Þ; (c) all faults of Salar Grande ðn ¼ 26Þ (note the confidence cone around the mean extension vector); and (d) faults of Mejillones Peninsula and Salar del Carmen areas ðn ¼ 39Þ: (e,f) Rose diagrams of poles of open cracks. (e) Salar Grande ðn ¼ 90Þ and (f) Salar del Carmen ðn ¼ 40Þ:

that there is some correlation. However, a detailed analysis of the data of the Salar Grande shows that there is a group of cracks with T-axes oriented N60 –70E, which differs from the orientation of T-axes related to faults.

8. Concluding remarks The kinematics of the three selected areas of northern Chile show that the Coastal Cordillera has undergone

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primarily bulk, near-surface, EW extensional deformation from the Pliocene to the Late Pleistocene. This strain field has been well documented along the Coastal Cordillera between the Rio Loa Canyon and Salar del Carmen area, including the Mejillones Peninsula, where deep parts of faults (50 m below the surface) are well documented. In the Salar Grande area, we find good evidence for near-surface EW extension. However, because the deep part of the structures in this area has not been well documented, our interpretation of the extent of the depth of the near-surface extension must be cautious. Another complication of this area is the occurrence of NW-striking dextral faults, which show a strain compatibility with normal faults. We interpret the occurrence of these dextral faults in terms of two working hypothesis: (1) in the framework of bulk trench-parallel extension, where dextral NW-striking faults are regarded as oblique transfer faults or (2) coastal parallel (N – S) shortening coupled with EW extension, where dextral faults formed contemporaneously with recently documented EW reverse faults. The exact origin of the N – S shortening cannot be identified using the available data. The horsetail fault geometry of the Salar Grande area, where dextral NW-striking faults link with the NS-trending Salar del Carmen Fault, could be inherited from the Mesozoic AFS. Furthermore, most faults of the Coastal Cordillera likely are Mesozoic structures that experienced reactivation during Neogene and Recent times. In the Mejillones Peninsula and Salar del Carmen, two documented roll-over structures indicate a trenchward tilting of the hangingwalls above faults whose dip angle decreases with depth. The 50– 100 m high, east-facing scarps of NS normal faults exposed in different parts of the Coastal Cordillera and the westward tilting of the hangingwall suggest that the faults have a listric geometry that dips eastward when viewed in section. In the Mejillones Peninsula, the increased amount of tilting with age indicates that the extensional slip was continuous from at least the Pliocene. In the Salar Grande area, the westward tilting of both the hangingwall and footwall of the NS-striking normal faults suggests that the fault’s architecture resembles a domino structure. The uplifted Late Pleistocene marine terraces (122 ^ 4 Ka and older) at the Mejillones Peninsula and along the coastal escarpment indicate that, in the longterm, coastal uplift is coeval with extensional deformation. However, the exact timing of the uplift relative to normal faulting is difficult to estimate from the available data. At the local scale, as in the Mejillones Peninsula, the hangingwalls and footwalls of normal faults seem to have experienced uplift through a deep-seated process. The normal faulting of the Mejillones Peninsula could be a consequence of the uplift of the central part of the peninsula, an idea supported by the along-strike variation of the slip magnitude of normal faults of the Mejillones Peninsula, whose tip point is located in the central part of

339

the isthmus. The observation that the highest uplift rate of the coastal area in northern Chile is recorded in the Mejillones Peninsula (Ortlieb et al., 1996a,b; Delouis et al., 1998), where the greatest Pleistocene vertical displacement has been documented, further supports the link between uplift and surface extension. In addition, the youngest faulting of the Coastal Cordillera is strongly correlated with those areas where marine terraces are broadly developed (Fig. 1). The mechanism that controls the near-surface extension of the emergent part of the forearc is probably related to the upward flexure of the forearc surface driven by plate convergence. The long wavelength, west flank of this flexure is subjected to gravitationally driven movement toward the trench that promotes extensional deformation. Extensional deformation results from a sort of large-scale buckle folding, in which normal faulting is concentrated in an outer arc above a neutral surface that separates it from an underlying inner arc that is undergoing contraction (Fig. 13). This hypothetical configuration may explain why the Caleta Herradura Fault is listric. The depth at which the fault becomes flat may represent the neutral surface, above which the crust is detached and moves toward the trench. 8.1. The relationship between long- and short-term deformations Comparing the coseismic westward displacements detected by GPS stations after the July 30 Antofagasta earthquake (Klotz et al., 1999), it is possible to demonstrate that there is a strong change in the magnitude of the coseismic vectors along the boundary between the Coastal Cordillera and the Central Depression (Salar del Carmen area). It indicates that a coseismic extensional strain accumulated along this boundary or in the Coastal Cordillera. This idea is in agreement with our observation that a larger cumulative extensional strain is concentrated in the Mejillones Peninsula and the Coastal Cordillera, whereas the inland part of the forearc is not deformed by extensional processes. The calculation of the strain magnitude between a GPS station located at the Coastal Cordillera relative to a GPS station in the middle part of the Central Depression (Klotz et al., 1999) gives a low value of , 0.0005%. The low strain magnitude is consistent with our observation that some open cracks were produced at relay ramps of the Salar del Carmen Fault during the last major earthquake in this area. Instantaneous stretching was probably triggered by great subduction earthquakes that caused local reactivation of the structures of the Coastal Cordillera. Furthermore, the fault striae on the Caleta Herradura, Mejillones, and Salar del Carmen Faults have a rake 73– 858 to the north, which indicates that the net separation along these faults has a small sinistral component. This is compatible with normal sinistral slip

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Fig. 13. (a) Geometry and kinematics of the outer Chilean forearc resulting from EW extensional deformation. In the Salar Grande area, extension is accommodated by NS-striking normal faults and NW-striking dextral faults. South of Salar Grande, the Coastal Cordillera is dominated by normal faults. (b) Interpretative EW schematic section showing the large-scale structural architecture of the Chilean continental margin. Directly beneath the Coastal Cordillera, strain and stress regimes are dominated by subduction of oceanic lithosphere that causes a compressive stress field, shortening, and uplift. The near-surface strain field is dominated by extension that occurs in a thin sheet of the upper crust. The boundary between the crust undergoing extension and contraction is a mechanical neutral surface. The normal faults curve into this surface and acquire a listric profile. The vertical magnitude of the extended area has been exaggerated.

along N – S faults that are reactivated by coseismic extension. Therefore, we propose that the coseismic extension is an additional process that promotes the trenchward movement of the forearc. 8.2. Alternative origins for open fractures Trench-perpendicular extension is also documented by NS-trending tensional cracks exposed along the Coastal Cordillera. However, oblique extension relative

to the normal component of the trench orientation has been recorded, as for the NNW-striking open cracks of the Salar Grande area, whose T-axes are oriented N60 –70E. Field observations suggest that some open fractures and normal faults have a closely related origin. However, because open fractures form only at low differential stresses, it is likely that they do not form exactly at the same time as their spatially associated normal faults. Furthermore, for open fractures to form in

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the crust, the minimum effective normal stress s3eff must be tensile. There are two end-member cases in which tensile stresses can be produced in the crust. The most common is driven by fluid pressures greater than the sum of the applied minimum principal stress and the tensile strength of the rock; that is, for an open fracture to form, Pf must equal s3 þ T: Alternatively, for a ‘dry’ rock, the tensile condition is only met when the minimum principal stress, s3 ; is tensile. There are two alternative ways to produce tensile stresses in the outer forearc of northern Chile. First, during the coseismic elastic rebound of the upper crust, the minimum subhorizontal principal stress becomes tensile. This triggers the formation of open cracks of low differential stress by ENE- to WSW-directed stretching of the crust. In this case, open cracks form perpendicular to the convergence direction. Such fractures may correspond to the NNW-striking open cracks documented in the depressed topography of Salar Grande, whose T-axes are oriented N60 –70E, nearly parallel to the present-day convergence direction between the Nazca and South American Plates. Second, normal faulting produces local, fast uplift that drastically reduces the confining pressure of the rocks, which leads to the local formation of tensile stresses by the gravitational force acting on the new, fault-formed, unconfined topography. Thus, local extension perpendicular to normal fault surfaces can form open cracks as secondary structures related to faulting.

Acknowledgements We thank the Universidad Cato´lica del Norte Research Fund for financial support through project DGICT 36000 and for supporting a research visit granted to GG at the Earth and Atmospheric Department of Cornell University. We thank Richard Allmendinger (Cornell University, USA), and Adrian Hatley (University of Aberdeen) who kindly revised an early version of this manuscript. Reynaldo Carrier (Universidad de Chile), ekkehard Scheuber (Freie Universita¨t Berlin Germany) and Jorge Skarmeta (Codelco, Chile) are thanked for the constructive criticisms. We also thank L. Jofre´ for help with the drafting. Fundacio´n Andes of Chile is currently funding our research in the region.

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