Field Trip Guides to the Backbone of the Americas in the Southern and Central Andes: Ridge Collision, Shallow Subduction, and Plateau Uplift (GSA Field Guide 13)

THF GEOLOGICAL SOCIETY OF AMERICA

Field Guide 13

Field llip Guides to the Backbone of the Americas in tlae Sotithem a11d Ccntra.IA11des: Ridge Collision, ShaDow Subduction, and Plateatl Uplift 2000 m (Biddle et al., 1986).

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Cretaceous The Cretaceous heralds the start of the “Andean” tectonic cycle as subduction along the western margin of South America began and was characterized by initial uplift of the Patagonian Cordillera, development of the Magallanes foreland basin, and emplacement of the majority of the Patagonian Batholith (Katz, 1972; Suárez and Pettigrew, 1976; Winslow, 1981, 1982). During the latest Jurassic and early Cretaceous, erosion of the Chon Aike volcanics provided clastic material for the fluvial and marginal marine sands that eventually formed the basal unit of the Magallanes basin, called the Springhill Formation (Riccardi, 1988). The Springhill is an important unit because it is the primary reservoir for hydrocarbon exploration in Magallanes basin. Subsequently, a thick sequence of early to middle Cretaceous black marine shales (the Rio Mayer and Rio Belgrano Formations), were deposited in an anoxic environment across the entire basin and provided the main source for hydrocarbons (Pittion and Gouadain, 1992). Uplift and deformation in the southern Patagonian Cordillera started in the middle to late Cretaceous and is marked by a peak in plutonic activity in the eastern margin of the Patagonian Batholith and by a change to coarse clastic sediments delivered to the Magallanes basin (e.g., the El Alamo, La Anita, Chorillo, Cerro Fortaleza, and Lago Sofia Formations; Winslow, 1982). From this point onward, the Magallanes basin developed as a true foreland basin as deformation progressed eastward toward the craton, producing the Patagonian fold-thrust belt (Winslow, 1981, 1982; Ramos, 1989; Kraemer, 1993; Klepeis, 1994; Coutand et al., 1999). Plate reorganization and/or an increase in plate convergence rates along the Patagonian margin are thought to have been responsible. Early Cenozoic In the Paleogene, continued uplift and deformation occurred within the southern Patagonian Cordillera (see reviews in Diraison et al., 2000; Ramos, 2005) as relatively rapid and steady convergence along the margin was maintained (Minster and Jordan, 1978; Pardo-Casas and Molnar, 1987; Gripp and Gordon, 1990). Paleogene uplift and deformation in the southern Patagonian Cordillera is marked by changing sedimentation patterns and deepening of the Magallanes basin westward toward the main cordillera with as much as 5000 m of sediment infill in the axial part of the basin (Biddle et al., 1986). Foreland basin sedimentation was dominated by continental and shallow marine sediments that are interpreted as synorogenic molasse deposits based on the presence of growth strata and prominent regional angular unconformities (Biddle et al., 1986; Malumián, 2002; Kraemer et al., 2002; Suárez et al., 2000). These sediments are youngest and record maximum deformation in the Fuegian Cordillera where the NNE convergence vector was more orthogonal to the margin (see Ramos, 2005). Unlike the Cretaceous, widespread subduction zone magmatism in the main cordillera is not observed (Ramos, 1982). Instead, large volumes of alkaline,

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OIB-like basalts erupt in the foreland basin and backarc regions, represented by the Posadas basalt. Peak magmatic activity for the Posadas basalts occurred at ca. 49 Ma, but they also show a broad southward younging age progression from north to south between 53 and 43 Ma (Ramos and Kay, 1992; Kay et al., 2002). Punctuating the overall steady convergence is the collision of the Farallon-Aluk spreading ridge system (Cande and Leslie, 1986). According to plate reconstructions by Cande and Leslie (1986), this ridge collision proceeded southward along the margin from ~42°S to the tip of Tierra del Fuego between ca. 60 and 40 Ma. Thus, there appears to be a close temporal and spatial correlation between the timing of ridge collision and deformation in the fold-and-thrust belt, cessation of arc magmatism, and the eruption of OIB-like alkaline basalts in the backarc (e.g., Ramos and Kay, 1992; Ramos, 2005). Late Cenozoic A new round of intense uplift and deformation occurred within the southern Patagonian Cordillera throughout the Neogene as the modern convergence geometry is established along the margin at ca. 25 Ma. The initial stage of Neogene uplift and deformation is marked by the deposition synorogenic molasse of the Rio Frias and Santa Cruz Formations in the early to middle Miocene that unconformably overly deformed Cretaceous and Paleogene sedimentary rocks. Superimposed on this relatively rapid convergence is the collision of the Chile Ridge spreading system beginning ca. 15 Ma (Cande and Leslie, 1986) and is thought to be responsible for the climax of Neogene uplift and deformation (Ramos, 2005). This ridge collision caused arc magmatism to shut down and shift eastward with the eruption of OIB-like alkaline basalts (Ramos and Kay, 1992; Gorring et al., 1997) and slab-melt adakites in the backarc (Kay et al., 1993; Ramos et al., 2004) and the emplacement of synorogenic granitoids on the eastern side of the main cordillera. A unique suite of forearc volcanics and granitoids was also emplaced along the western coastal belt on the Taitao Peninsula (Mpodozis et al., 1985; Forsythe et al., 1986; Lagabrielle et al., 1994; Le Moigne et al., 1996; Guivel et al., 1999; Lagabrielle et al., 2000). Neogene topographic uplift is also thought to have contributed to intense glaciation of the southern Patagonian Cordillera, starting around 6 Ma (Mercer, 1976) and continuing into the Holocene (Ivins and James, 1999). LATE CENOZOIC TECTONIC FRAMEWORK The current tectonic framework of the southern Andean Cordillera involves a relatively complex interaction between the oceanic Nazca, Antarctic, and Scotia plates and the continental South American plate (Fig. 1). The Nazca plate subducts rapidly beneath the South American plate at a relative velocity of 9 cm/yr, whereas the Antarctic plate subducts more slowly at 2 cm/yr. The Nazca and Antarctic plates are separated by the Chile Ridge system. At the southernmost tip of the Andes,

the Scotia and South American plates form a large-scale, leftlateral transcurrent boundary. The current plate motion vectors and relative convergence rates were established ca. 25 Ma, when the Nazca plate vector changed from highly oblique (010°E) to approximately orthogonal (080°E) with respect to the continental margin of South America (Minster and Jordan, 1978; PardoCasas and Molnar, 1987; Gripp and Gordon, 1990). Beginning at ca. 14–15 Ma, the Chile ridge system collided with the southernmost tip of the Patagonian Andes, in the western part of Tierra del Fuego (Cande and Leslie, 1986). The Chile Triple Junction (the triple point between Nazca, Antarctica, and South America) has since migrated northward along the margin in a series of ridge collision events to its present location near the Taitao Peninsula at 46.5°S (Cande and Leslie, 1986). Thus, since the middle Miocene, the tectonics along the margin of the southern Patagonian Cordillera south of Chile Triple Junction has changed from rapid (9 cm/yr), slightly oblique (075°E) convergence associated with subduction of the Nazca plate to slow (2 cm/yr), orthogonal (090°E) convergence associated with Antarctic plate subduction. Because the Chile Ridge system is segmented with individual ridge axes oriented NNW-SSE, the ridge collision is only slightly oblique to the margin. This relatively simple collision geometry coupled with the rapid (7 cm/yr) westward absolute plate motion vector of the South American plate is likely responsible for the complete subduction of the Chile Ridge system without any trace of internal deformation in either the Nazca or Antarctic plates. CONSEQUENCES OF RIDGE COLLISION The primary, large-scale geodynamic consequence of the ridge-trench collision is the formation of asthenospheric slab windows beneath the southern Patagonian Cordillera (Cande and Leslie, 1986; Ramos and Kay, 1992; Gorring et al., 1997) (see Fig. 2). Slab windows form because of the large differential convergence velocities (~7 cm/yr) between the Nazca and Antarctic plates (Gorring et al., 1997). In theory, the opening of slab window allows relatively hot, asthenospheric mantle to flow upward between plates (e.g., Thorkelson, 1996), and this process has been linked to profound effects on the late Cenozoic magmatic and deformational history of the southern Patagonian Cordillera (Ramos and Kay, 1992; Gorring et al., 1997; Ramos, 2005). The unique geodynamics of ridge collision and slab window formation, in theory, should have profound, observable effects on the geologic evolution of a mountain belt. In the southern Patagonian Cordillera, the following features are thought to be related (either directly or indirectly) to the late Cenozoic ridge-trench collision: • Ophiolite emplacement (Mpodozis et al., 1985; Guivel et al., 1999); • Forearc subduction erosion (Bourgois et al., 1996); • Anomalous forearc felsic and MORB-like magmatism (Mpodozis et al., 1985; Forsythe et al., 1986; Lagabrielle et al., 1994; Le Moigne et al., 1996);

Figure 2. Schematic cross sections (no vertical exaggeration) showing the Patagonian slab window model (Gorring et al., 1997), highlighting mantle source regions and petrogenetic processes involved in the genesis of Neogene slab window lavas erupted northeast of where a Chile Ridge segment collided with the Chile Trench at ca. 12 Ma. Abbreviations: OIB—oceanic island basalt; SSVZ—southern Southern Volcanic Zone.

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• Formation of a gap in the Quaternary volcanic arc (Stern et al., 1990; Mpodozis et al., 1985); • Topographic uplift and reactivation of deformation in the Patagonian fold-and-thrust belt (Ramos, 1989; Coutand et al., 1999; Kraemer et al., 2002; Ramos, 2005); and • Adakitic and OIB-like mafic magmatism in the backarc (Ramos and Kay, 1992; Gorring et al., 1997). This field trip will focus on the last two items on the bulleted list above, and these items are described in further detail below. PATAGONIAN FOLD-AND-THRUST BELT Basic Structure The Patagonian fold-and-thrust belt is a classic foreland fold-thrust belt that extends for ~1000 km along the eastern foothills of the southern Andes between 46° and 55°S along the southwestern margin of the Magallanes basin (Winslow, 1982; Ramos, 1989; Klepeis, 1994; Kraemer, 2003) (Fig. 1). The belt is ~40–100 km wide and can be generally split into two along-strike segments, an eastern foreland zone and a western hinterland zone (Ramos, 1989; Kraemer, 1993) (Fig. 3). In the sector north of 51°S, deformation in the eastern foreland zone is characterized by gentle, kilometer-scale folding and thinskinned, west-verging backthrusts in Cretaceous and Tertiary sedimentary rocks (Ramos, 1989; Kraemer, 1993; Coutand et al., 1999). Deformation gradually increases westward toward the hinterland, where deformation is characterized by mostly thick-skinned, east-verging imbricate thrust sheets that uplift late Paleozoic basement and Mesozoic volcanic rocks (Fig. 4). A triangle zone marks the transition between the foreland and hinterland zones, which is particularly well developed north of Lago San Martin (~49°S; Ramos, 1989) (Figs. 3 and 4). Major décollements are recognized to occur near the base of the Rio Mayer Formation and within the late Paleozoic metasedimentary basement where much of the shortening is accommodated. Fault kinematic analysis and the N-S to NNW trend of folds and thrusts both indicate dominant E-W compression with a component of right-lateral wrenching along strike in this sector of the Patagonian fold-and-thrust belt (Coutand et al., 1999). Topography and Timing of Deformation in Relation to Late Cenozoic Ridge Collision The timing of deformation in the southern Patagonian Cordillera and the development of the Patagonian fold-and-thrust belt is broadly constrained by important changes in sedimentation and the presence of unconformities in the Cretaceous and Tertiary section. There is general consensus that the initial formation of the fold-and-thrust belt started during the middle to late Cretaceous and was followed by major contractional events that took place during the latest Cretaceous, Eocene, and Miocene times (e.g., Ramos, 1989; Suárez et al., 2000). These major deforma-

tional events have been linked to periods of rapid orthogonal convergence at ca. 80 Ma, ca. 50–40 Ma, and 25–10 Ma (e.g., Suárez et al., 2000), but were also enhanced by ridge collisions events during these times (e.g., Ramos and Kay, 1992; Ramos, 2005). With respect to the late Cenozoic event, the beginning of deformation is constrained by synorogenic molasse deposits of the Rio Frias and Santa Cruz Formations that contain interbedded ash layers with maximum Ar/Ar ages of ca. 19 Ma (Feagle et al., 1995). An angular unconformity exists between the Santa Cruz Formation and the overlying main plateau basalts, the oldest of which are ca. 14–12 Ma (Gorring et al., 1997), and constrains the minimum age of significant foreland basin sedimentation. Additional evidence for latest Oligocene to mid-Miocene deformation comes from apatite fission track data, which suggest that rapid uplift and denudation started ca. 30–23 Ma along the Pacific coast and subsequently migrated 200 km eastward until ca. 12–8 Ma (Thomson et al., 2001). New age data and structural information from the Torres del Paine region (51°S) suggest significant late Oligocene to mid-Miocene compressional deformation constrained by the deformed “external gabbros” dated at ca. 30 Ma and the undeformed Torres del Paine pluton with a minimum age of ca. 12 Ma (Altenburger et al., 2003). Oxygen isotope data from paleosols from the Santa Cruz Formation indicate that the present-day orographic rain shadow across the southern Patagonian Cordillera was established between ca. 17 and 14 Ma and can be attributed to rapid topographic uplift of >1 km (Blisniuk et al., 2005). The above data clearly indicate that uplift and deformation was well under way prior to collision of the Chile ridge system and is linked to more orthogonal and increased convergence rates at ca. 25 Ma (e.g., Ramos, 1989; Suárez et al., 2000; Thomson et al., 2001). However, there is also evidence that final uplift and deformation in the eastern main cordillera and the Patagonian foldthrust belt is linked to the late Cenozoic ridge collision along the southern Patagonian margin. Ramos and Kay (1992) and Ramos (2005) pointed out the drastic change in the topography and style of deformation that occurs at the latitude of the modern Chile Triple Junction (46.5°S). There is an abrupt uplift of >2000 m of elevation along the crest of the Patagonian Cordillera from north to south at 46.5°S (Fig. 5). To the north, the average elevation of the highest peaks of the Patagonian Cordillera is ~2000 m, whereas to the south of the Chile Triple Junction, average elevations increase suddenly to >4000 m (Cerro San Valentin, 4078 m) and remain relatively high at >3000 m (Cerro San Lorenzo, 3706 m; Cerro Fitz Roy, 3405 m; Cerro Paine Grande, 3050 m; among others) until 53°S, where once again maximum average peak heights are ~2000 m. This difference in topography is spatially correlated with a significant difference in the style of deformation north and south of the Chile Triple Junction and the development of the southern Patagonian fold-thrust belt (Ramos, 2005). South of the Chile Triple Junction, uplift of the southern Patagonian Cordillera is accomplished by crustal stacking that involves substantial amounts of shortening taken up by both thin-skinned (foreland) and thick-skinned deformation (main cordillera).

Figure 3. Major structures of the Patagonian fold-and-thrust belt. Figure from Ramos (1989).

Figure 4 (on this and following page). Structural cross sections of the Patagonian fold-and-thrust belt. Lines of section are in Figure 3. Figure from Ramos (1989).

Figure 4 (continued ).

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Figure 5. North-south topographic section of the Patagonian Andes in which maximum elevation is indicated at each latitude. Present Chile Ridge collision is taking place at 46.5°S. Figure from Ramos (2005).

North of the Chile Triple Junction, there is only a modest amount of shortening, with crustal stacking taking place primarily through mild tectonic inversion of Mesozoic normal faults coupled with significant dextral transpressional deformation taken up on the Liquine-Ofqui Fault (Fig. 6). Further evidence for Late Cenozoic ridge collision-related uplift and deformation comes from isotopic ages (mostly K/Ar and 40Ar/39Ar) of granitoid plutons from the eastern edge of the main cordillera. Ages range between 18 and 3 Ma and include the following: • Cerro Fitz Roy pluton (18 ± 3 Ma; Nullo et al., 1978); • Torres del Paine pluton (13 ± 1 and 12 ± 2 Ma; Halpem, 1973; Michael, 1983); • Paso de las Llaves pluton (ca. 10 ± 0.5 Ma; Petford and Turner, 1996; Pankhurst et al., 1999; Suárez and de la Cruz, 2001; Thomson et al., 2001); • Cerro San Lorenzo Pluton (ca. 6.5 ± 0.5 Ma; Welkner, 1999; Suárez and de la Cruz, 2001); and • Rio de las Nieves pluton (3.2 ± 0.4 Ma; Morata et al., 2002). The mid-Miocene to Pliocene ages coupled with the present elevation of >2000–4000 m of the Cerro Fitz Roy, Cerro San Lorenzo, and Torres del Paine plutons clearly require significant post-middle to late Miocene exhumation and erosion of cover rocks (e.g., Skarmeta and Castelli, 1997; Suárez et al., 2000; Ramos, 2005). Coutand et al. (1999) cited evidence for Pliocene shortening in Patagonian fold-thrust belt along the north shore of Lago Viedma (49.5°S) that includes gentle tilting of Pliocene plateau basalts and feeder dikes that cut Early Cretaceous sediments that are offset with top to the east (reverse) sense of motion. In the Lago Buenos Aires region (~46.5°S), Lagabrielle et al. (2004) cited geomorphic evidence for post– late-Miocene uplift, including uplift and dissection of relict late Miocene-Pliocene paleosurfaces, stream capture, and transpressional strike-slip faults that cut late Miocene plateau basalts.

BACKARC MAGMATISM RELATED TO RIDGE COLLISION Perhaps the most unequivocal affects of ridge collision in the southern Patagonian Cordillera are the backarc magmatic affects. Suites of distinctive igneous rocks are well characterized geochemically and are well constrained by radiometric dating and thus can be correlated in both and time and space with the sequential collision of segments of the Chile Ridge along the southern Patagonian margin since the middle Miocene (Gorring et al., 1997; Ramos et al., 2004). Adakites The term “adakite” was coined by Defant and Drummond (1990) for a geochemically distinctive type of silicic volcanic rocks from Adak Island in the Aleutians. These rocks were originally discovered and interpreted by Kay (1978) as being generated by partial melting of oceanic crust (e.g., “slab melting”). Since 1990, the term adakite has been applied (controversially) to a variety of volcanic rocks with “adakitic” geochemical characteristics that may have formed from distinctly different processes other than direct slab melting, namely forearc subduction erosion and partial melting of thickened mafic lower continental crust (see Kay and Kay, 2002). Thus, the origin of many adakites via direct slab melting has been vigorously debated (see Yogodzinski et al., 2001). Perhaps the best remaining candidates for a slab-melt origin are those from southern Patagonia that erupted in the backarc region east of the modern volcanic arc gap between the Austral Volcanic Zone and Southern Volcanic Zone and where ridge subduction has occurred over the past ca. 12 Ma (Kay et al., 1993; Ramos et al., 2004). Adakites from three separate localities have been recognized: these are the Chaltén (49.2°S), Puesto Nuevo (48.6°S), and Cerro Pampa (47.6°S) adakites. Outcrops at all

Ridge-Trench Collision—The Southern Patagonian Cordillera

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Figure 6. Schematic cross sections highlighting major structural differences of the Patagonian Cordillera north (A) and south (B) of the present Chile Triple Junction. Figure from Ramos (2005).

three localities are relatively small (~100–200 m diameter), pluglike, subvolcanic (?) bodies of porphyritic dacite with large (up to 4–5 cm long), acicular phenocrysts of hornblende ± plagioclase. Convincing geochemical evidence for a slab-melt origin for these hornblende dacites comes from high Sr (1330–2300 ppm), Cr (80–100 ppm), and Ni (40–75) at 63%–68% SiO2, MORB-like 87 Sr/86Sr (0.7028–0.7033), 143Nd/144Nd >0.51289, and steep rare earth element (REE) patterns (La/Yb = 28–38, heavy rare earth element [HREE]-depleted) (Kay et al., 1993; Ramos et al., 2004) (Table 1). New 40Ar/39Ar laser ablation dates on hornblende for Chaltén, Puesto Nuevo, and Cerro Pampa adakites show a systematic northward decrease in age from ca. 14.5, ca. 13.1, to ca. 11.5 Ma, respectively (Ramos et al., 2004). The geochemistry and timing of these adakites is consistent with partial melting of the young, hot trailing edge of the Nazca plate associated with the ca. 12 Ma ridge collision event that preceded slab window opening and the eruption of extensive OIB-like basalts (Fig. 2).

OIB-Like Slab Window Basalts Large volumes of mafic slab window magmas erupted over vast areas of the southern Patagonian backarc southeast of the modern Chile Triple Junction following a series of ridge collisions along the Chile Trench during the mid- to late Miocene (Gorring et al., 1997). Slab window lavas are most abundant between 46.5° and 49.5°S, northeast of two ridge segments that collided at ca. 12 Ma and ca. 6 Ma and are located 100–400 km east of the volcanic arc gap between the Southern Volcanic Zone and Austral Volcanic Zone (Fig. 1). K/Ar and 40Ar/39Ar ages (Ramos and Kay, 1992; Gorring et al., 1997) suggest two periods of magmatism: (1) an older (12–5 Ma) voluminous, tholeiitic (48%–55% SiO2; 4%–5% Na2O + K2O) main-plateau sequence and (2) a younger (7 to 50 km) with seismic data indicating thicknesses as much as 65–80 km in the central plateau (Yuan et al., 2002; McGlashan et al., 2008). The highest regions correspond with sites of Neogene volcanic activity. Uplift of the plateau is considered to have largely occurred in the Neogene with the principal cause being crustal thickening in response to crustal shortening along with limited magmatic addition (e.g., Allmendinger et al., 1997). Delamination of the lower continental crust and lithosphere and the resulting thermal conditions have contributed to the high elevation (e.g., Kay and Kay 1993; Kay et al., 1994a). General overviews on the plateau have been presented in Allmendinger et al. (1997), Kay et al. (1999, 2004), Beck and Zandt (2002), Oncken et al. (2006), and references in those papers. The central Andes are characterized by a dominantly compressional Neogene stress regime not related to continental collision. Contractional deformational belts of Neogene age border the eastern side of the plateau (Fig. O.1; see Allmendinger et al., 1997). From north to south, they include the Subandean and Eastern Cordilleran fold-thrust belts, the Santa Barbara belt where shortening is accommodated by inversion of Cretaceous normal faults, and the high-angle, reverse-faulted Sierras

Pampeanas. The amount of shortening is variable, and exact amounts are widely debated (e.g., Kley and Monaldi, 1998; Kley et al., 1999). The subducting Nazca plate beneath the central Andean plateau is Oligocene to Eocene in age. A feature of this slab is its relatively shallow angle (~30°) compared to other regions around the circum-Pacific where subduction angles are rarely less than 45°. To the north and south of the plateau lie shallower subducting segments under the volcanically quiescent Peruvian and Chilean flat-slab regions (Fig. O.1; Isacks 1988; Cahill and Isacks, 1992). The southward transition to the Chilean flat slab is relatively smooth, whereas that with the Peruvian flatslab segment is abrupt. Flattening is expressed as a bench that develops in the seismic zone between 90 and 135 km (Fig. O.1; Cahill and Isacks, 1992). Gephart (1994) emphasized the degree of bilateral symmetry in the shape of the seismic zone and topography of the land surface from 33°S to 5°S latitude. Isacks (1988) argued that the modern geometry of the subducting Nazca plate and uplift of the central Andean plateau are the result of the oroclinal bend in the South American plate overriding the subducting Nazca plate. Models for the evolution of the plateau call for a “collision” between the shallowly dipping Nazca plate and the overriding South American plate. Recent papers have emphasized the importance of the westward drift of South America (Sobolev and Babeyko, 2005; Oncken et al., 2006) as originally proposed by Silver et al. (1998). The central Andean plateau can be divided into the Puna (~22° to 27°S) and the Altiplano (~22° to 15°S) as discussed by Turner (1970). Some important differences are shown in Figures O.2 and O.3. As summarized by Whitman et al. (1996), Allmendinger et al. (1997), and Yuan et al. (2002), the Puna differs from the Altiplano in that: (a) the basement is generally younger and has a larger component of Paleozoic magmatic rocks; (b) small discontinuous, diachronous basins replace the large Altiplano basin; (c) widespread Miocene to Quaternary volcanic rocks erupted

Neogene evolution: central Andean Puna Plateau and southern-central volcanic zone

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Figure O.1. (A) Map of the Puna-Altiplano region of the central Andes. Area above 3700 m in elevation is in red; fold-and-thrust belts are in green; Sierras Pampeanas uplifts are in gray. Contours to the Wadati Benioff zone are labeled at 50 km intervals (Cahill and Isacks, 1992). CVZ—Central Volcanic Zone; SVZ—Southern Volcanic Zone. (B) Distribution of slab earthquakes (dots) from Cahill and Isacks (1992). The region in red shows the area of mafic lavas (Kay et al., 1994a); circle encloses region of lavas with intraplate chemistry. Purple area has low seismic Q following Whitman et al. (1992, 1996).

in broad northwest-trending chains separated by regions of basement outcrop (Fig. O.2); (d) average elevations are higher (Figs. O.2 and O.3); and (e) seismic, topographic, and gravity data support a thinner lithosphere and crust (Figs. O.3–O.6). Geophysical characteristics of the northern Puna and Altiplano are better known due to the seismic studies of ANCORP (2003) and the Arizona group (e.g., Beck and Zandt, 2002). The ANCORP seismic profile near 22°S is shown in Figure O.5. A seismic array was installed in the southern Puna by a Cornell-MissouriPotsdam group in 2007. The Puna Plateau The modern Puna plateau shows some important north to south differences as initially suggested by Alonso et al. (1984a).

Among the contrasts are a higher average topography (Isacks, 1988) and a thinner lithosphere in the south. Evidence for the thinner lithosphere comes from asthenospheric shear wave attenuation, lower seismic Q, and effective elastic thickness (Whitman et al., 1992, 1996). Evidence for a thinning lithosphere at the transition between the northern and southern Puna comes from seismic tomography (Schurr et al., 2003, 2006) and Quaternary shoshonitic lavas erupted along faults (Coira and Kay, 1993; Kay et al., 1994a). Gravity data support a region of distinctively thin lithosphere and crust near 25°S (Tassara et al., 2006) and are in accord with seismic data indicating crustal thicknesses near 42–49 km (Fig. O.4). Crustal thicknesses from 22°S to 25°S range from ~50 to 68 km; thicknesses are unavailable farther south (Fig. O.4). The southern Puna also differs from the northern Puna in having normal and strike-slip faults (e.g., Marrett et al.,

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Figure O.2. Satellite Radar Topography Mission (SRTM) image of the central Andean plateau showing the Altiplano basin bordered by the Eastern and Western Cordillera in the Altiplano, the Altiplano-Puna Volcanic Complex ignimbrites (APVC of de Silva, 1989), the isolated basins of the southern Puna, and the Cerro Galan caldera. The field trip is within the black box.

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Figure O.3. Lithospheric cross section illustrating north to south differences across the Puna-Altiplano in topography, lithospheric thickness, and subducting slab geometry. Figure is modified from Whitman et al. (1996).

1994) associated with 1600 km3 La Pacana caldera at ca. 4 Ma (e.g., Lindsay et al., 2001) and the ~1500 km3 Puripicar ignimbrite erupted at ca. 4.2 Ma (Barquero-Molino, 2003; de Silva and Gosnold, 2007). By the latest Pliocene, nearly all magmatic activity was concentrated near the modern Central Volcanic Zone arc front. Thrusting continued in the Subandean Belt with out-of-sequence thrusting occurring after 4.5 Ma (Echavarria et al., 2003). Southern Puna History (Figs. O.7 and O.8) From 27 to 19 Ma, widespread ignimbrite and dacitic dome complexes were erupting in the frontal arc west of the Puna. Backarc activity was limited to mafic volcanism just east of the arc near 26°S (Segerstrom basalt; Kay et al., 1999) and small dacitic to rhyodacitic complexes in the Puna (Coira et al., 1993). This picture changed between ca. 20 and 16 Ma when small stratovolcanoes and minor ignimbrites began erupting in the backarc and contractional deformation took place east of the Maricunga Belt (Mpodozis and Clavero, 2002). By the middle Miocene, 16–12 Ma andesitic stratovolcanoes were erupting in the Maricunga Belt arc, and volcanism initiated at ca. 15–14 Ma at the long-lived backarc stratovolcanic complexes like Beltran, Antofalla, and Tebenquicho (e.g., Coira

et al., 1993; Kraemer et al., 1999; Richards et al., 2006). Ignimbrites erupted in the transitional region between the northern and southern Puna near 24°S (Petrinovic, 1999; Petrinovic et al., 1999). The magmatic style changed at 11–7 Ma as volcanism in the Maricunga Belt arc became concentrated in the Copiapó dacitic ignimbrite-dome complex (Mpodozis et al., 1995; Kay et al., 1994b) and local 11–10 Ma ignimbrites in the backarc preceded andesitic lavas and dome complex formation in the long-lived Puna stratovolcanic complexes. Contractional deformation continued in the southern Puna, and the Pampean ranges east of the Puna began to uplift at ca. 9–8 Ma. By ca. 6–5 Ma, volcanism had terminated in the Maricunga Belt arc as frontal arc activity shifted eastward toward the Central Volcanic Zone arc (Mpodozis et al., 1996). A major change occurred in the Puna at ca. 6.7 Ma as mafic lavas erupted along normal and strike-slip faults (e.g., Kay et al., 1999) and the Cerro Galán ignimbrite eruption began (Sparks et al., 1985). By ca. 4 Ma, ignimbrites were erupting in both the arc (Laguna Verde, Amarga, Vallecito; Mpodozis et al., 1996; Siebel et al., 2001) and backarc (Real Grande ignimbrite at Cerro Galan; Sparks et al., 1985). The arc front was stabilized in the Central Volcanic Zone (Fig. O.8.; Mpodozis et al., 1996) by 3–2 Ma and the ~1000 km3 Cerro Galán ignimbrite erupted at ca. 2.2 Ma

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(Sparks et al., 1985). During this time, intraplate-like mafic lavas erupted over the modern gap in intermediate depth slab seismicity (intraplate region in Figs. O.1B and O.8) and arclike magmas erupted to the north and south (Kay et al., 1994a, 1999). Since 2 Ma, silicic magmas have erupted from the Cerro Blanco caldera, and Puna mafic volcanism has continued (Siebel et al., 2001; Kay et al., 2006; Risse et al., 2008). Models for the Neogene Magmatic-Tectonic Evolution of the Northern and Southern Puna Northern Puna: Shallowing and Steepening of the Subduction Zone (Fig. O.9) The model for the evolution of the northern Puna in Figure O.9 involves a transition from a very shallow, late Oligocene to early Miocene subduction zone to a moderately steep subduction zone (Coira et al., 1993; Kay et al., 1999). Periodic delamination of the lower crust and lithosphere could have accompanied slab steepening (e.g., Beck and Zandt, 2002). Stage 1 at 26–14 Ma: The virtual lull in volcanism, widespread contractional deformation, and basin formation across the arc and backarc are consistent with a shallowly dipping subduction zone like that under the modern Chilean flat-slab region. A similar Oligocene flat slab has been suggested by James and Sacks (1999) to explain geophysical and geological observations under the southern Altiplano. Stage 2 at 14–6 Ma: Steepening of the subduction zone after ca. 14 Ma is consistent with small backarc dacitic eruptions preceding widespread, voluminous late Miocene ignimbrite eruptions. The transfer of contractional deformation into the Subandean Belt and important uplift of the plateau after 10 Ma fit with the model of Isacks (1988) in which upper crustal shortening is compensated by ductile lower crust under the plateau at the time of uplift. Accumulation and fractionation of lower crustal melts at the brittle-ductile transition near 20 km depth is in accord with magma chambers inferred near that depth from geophysical data by Zandt et al. (2003) and on the ANCORP (2003) profile (Fig. O.5). The horizontal compressional failure of melt-weakened crust can explain the transfer of these magmas to the shallow crustal chambers from which they are inferred to erupt (e.g., Lindsay et al., 2001). The major ignimbrite eruptions are also potentially linked to the crustal failure that produced the Subandean thrusts. Delamination of the crust and mantle lithosphere under the northern Puna and southern Altiplano (Yuan et al., 2002; Beck and Zandt, 2002; Garzione et al., 2006) could have enhanced crustal melting by intrusion of mantle melts as argued by Kay and Kay (1993) for Cerro Galán. However, in detail, this delamination needs to differ from that under the southern Puna because the mafic lavas and mixed extensional and/or strike slip and/or contractional fault system found in the southern Puna are essentially absent. Stage 3 at 6–3.8 Ma: Further steepening of the subducting slab is consistent with eruption of giant ignimbrites near the modern arc front and continuing contraction in the Subandean Belt.

Figure O.9. Lithospheric-scale sections to explain the magmatic and deformational history of the northern Puna modified from Kay et al. (1999) and Kay and Mpodozis (2001). CVZ—Central Volcanic Zone.

Stage 4 at 25) are consistent with the magmas equilibrating with residual garnet-bearing lower crust. The glassy character suggests that the flows erupted as hot, degassed magmas that ascended rapidly along faults. A glassy andesite on the northeastern margin of the Laguna Caro depression has a whole-rock 40Ar/39Ar age of 4.6 ± 0.5 Ma (Risse et al., 2008). Stop 3-5: View of Salar de Antofalla, Antofalla, and Tebenquicho Stratovolcanoes Directions: Continue west on dirt track to 25°34′58″S; 67°30′54″W; 3928 m asl. Stop at top of ridge where road descends into the Salar de Antofalla. A map of the area is in Figure 3.4.

C° Tebenquicho Grande

68°00′ W

N

C° Onas

C° Los Patos

25°30′ S

C° Conito Antofalla

l l a

Antofalla

5

f a

Vn. Antofalla

C° Lila

Potrero Grande

a

Campo del Volcán

l a

r

6

S

C° de la Aguada

A

Puesto Cuevas

C° Cajeros

n

t o

C° Bayo

Vega Botijuela

Quaternary

Clastic-pyroclastic sequences Eocene-Pliocene

Undifferenciated sediments

Upper Paleozoic Mesozoic Lower Paleozoic Upper Precambrian

5 km

0.1-0.5 Ma

Lavas and scoria cones Basalt/basaltic andesite

1.5-1.2 Ma

Lavas and scoria cones Basalt/basaltic andesite

2.3 Ma

Cuevas/Botijuela rhyolite vitrophyre

Fault Anticline Syncline

Lava flows and scoria cones Basaltic andesite to andesite 7-4 Ma Reverse fault Andesite-dacite composite volcanoes: d) Lava flows 9-8 Ma d 12-13 Ma c) Andesite-dacite lava, minor basaltic andesite lava c 10.8 Ma b) Dacite-rhyolite ignimbrite A1 b a2 14-10 Ma A) Central volcanic complex strongly eroded: 1) Basalt - andesite 2) Dacite-andesite and rhyolite lava domes Sedimentary rocks Metasediments, metavolcanics, amphibolites, schist-gneiss

Figure 3.4. Map by B. Coira of the Antofalla volcanic complex and region near Stop 3-5.

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The Antofalla stratovolcano on the western margin of the Salar de Antofalla to the west is a voluminous and long-lived 6400 m volcanic complex with a diameter of 35 km. The center erupted in three main stages. The first stage produced: (1) basaltic andesitic lava cones like Cerro de la Aguada (12.9 ± 0.8 Ma; Kraemer et al., 1999), Cajero (13.0 ± 0.5 Ma; Schnurr, 2001) and Antofalla (13 Ma; Coira and Pezzutti, 1976); (2) dacitic lava domes like Cerro Lila (10–9 Ma; Coira and Pezzutti, 1976); (3) dacitic to rhyolitic ignimbrites (10.9–9.6 Ma; Kraemer et al., 1999); (4) andesitic to dacitic lavas (9–8 Ma; Coira and Pezzutti, 1976); and (5) minor basaltic andesitic flows erupted along WNW-ESE and north-south structures. In the next stage between 7 and 4 Ma, basaltic andesites and minor basalt flows erupted on the margins of the complex and from flank vents on the older centers. Basaltic, basaltic andesite, and andesitic flows also erupted from structurally controlled vents on the eastern and southern margins of the Salar de Antofalla. The last stage produced sparse basaltic andesitic lavas and small rhyolitic lava domes like those at Las Cuevas and Cerro Botijuela (2.5–2.3 Ma; Schnurr, 2001; Richards et al., 2006). Pleistocene (1.5–1.2 and 0.1–0.5) andesitic to basaltic scoria cones and lava flows (Marrett et al., 1994; Kay et al., 1997) also erupted on the margins of the Salar de Antofalla. The Tebenquicho stratovolcanic complex to the north is dominated by andesitic to dacitic lava flows that erupted between 14 and 6 Ma (Kraemer et al. 1999). The Antofalla and Tebenquicho volcanic complexes unconformably overlie thick sequences of Cenozoic sediments that record the tectonic-sedimentary evolution of the region. Figure 3.5 from Carrapa et al. (2005) provides a summary in the Salar de Antofalla region. Sedimentation began in the latest Eocene with deposition of the clastic foreland basin deposits in the late Eocene to Oligocene Quiñoas Formation that were derived from the west as a result of the Incaic deformation (Jordan and Alonso, 1987; Kraemer et al., 1999; Voss, 2002). An arid climate was established at this time. Late Oligocene thick-skinned contractional deformation (D1 in Figure 3.5; Adelmann and Görler, 1998) triggered syntectonic deposition of the coarse-grained late Oligocene to early Miocene Chacras Formation alluvial fan deposits (Kraemer et al., 1999). These fan sediments were largely derived from the Sierra de Calalaste region farther south (Carrapa et al., 2005), which will be visited on Day 4. Uplift associated with Oligocene deformation led to reorganization of the depositional systems in the Salar de Antofalla area. Renewed east-west to WNW-ESE shortening in the early Miocene (D2 in Fig. 3.5, ca. 20–17 Ma) reactivated the Paleogene west-vergent fault system (Adelmann and Görler, 1998). During this time, the Salar de Antofalla region was separated into small intra-arc depocenters in which the Potrero Grande Formation alluvial fan and fluvial sediments accumulated (Kraemer et al., 1999; Voss, 2002). The Miocene west-vergent thrusts affected Lower Paleozoic, Permian, and Tertiary rocks. Younger Miocene shortening (D3 in Fig. 3.5) produced east- and west-vergent basement thrusts that tilted the Potrero Grande Formation alluvial fans and generated further deposition. The middle Miocene to Pliocene Juncalito Formation

of Kraemer et al. (1999) was deposited at this time. The thick evaporate deposits in these sequences show that the Salar de Antofalla basin was internally drained by the late Miocene. A mixed Pliocene stress regime (D4 in Fig. 3.5) produced both contractional and local strike-slip deformation. The present narrow and elongate shape of the Salar de Antofalla basin reflects this deformation as well as Quaternary erosional processes. Unlike other southern Puna salars, evaporates in the Salar de Antofalla are largely halite rather than borate deposits. Seismological studies near 25.5°S by the ANCORP group provide information on the crust beneath the Salar de Antofalla (Heit, 2005). A high-velocity anomaly coincides with a portion of the central Andean gravity high that Götze and Krause (2002) associated with Lower Paleozoic ultrabasic rocks in the basement. The sides of the Antofalla depression are flanked by low-velocity anomalies. The anomaly to the west extends under the main Central Volcanic Zone arc. The anomaly to the east at ~67°W is along strike with the Cerro Galán caldera. The shape and extent of the narrow northeast-trending Salar de Antofalla and the topographic difference between the salar surface (3400 m) and the

Quaternary Escondida Pliocene

10

Fm.

D4

Juncalito Fm. D3 Potrero Grande Fm.

20

D2

Chacras Fm. D1 30 Quinoas Fm. Incaic 40 no sedimentation

Figure 3.5. Summary of the deformational and stratigraphic development of the Salar de Antofalla and surrounding region as compiled by Carrapa et al. (2005) based on the synthesis of Kraemer et al. (1999), Adelmann (2001), and Voss (2002). Dashed lines between formational names indicate important regional scale unconformities.

Neogene evolution: central Andean Puna Plateau and southern-central volcanic zone flanks (~4000 m) appears to reflect a deeper structure bounded by these velocity anomalies. In November of 1973, a magnitude 5.8 earthquake occurred at a depth of 6 km beneath the region (Chinn and Isacks, 1983). The kinematics of the focal mechanism indicate north-south–dipping, 30° extension, and east-west horizontal shortening. One of the nodal planes strikes parallel to the major axis of the Salar de Antofalla. Stop 3-6: Faults and Mafic Lavas in the Vega de los Colorados Directions: Return to Vega de los Colorados; stop at 25°35′34″S; 67°30′54″W; 3928 m asl (Figs. 3.2 and 3.6).

14–10 Ma lavas

7–4 Ma mafic lavas

Tertiary clastic sequence

Pc/Pal

149

Examples of the strike-slip faults in the southern Puna can be observed in the Vega de los Colorados. These faults have been studied by Marrett et al. (1994) and Kraemer et al. (1999). The main branch of the regional-scale Acazoque fault and a branching synthetic strike-slip fault are well seen. The Acazoque fault strikes NNE-SSW and dips southeast. The fault is marked over much of its length by a conspicuous scarp along which basaltic andesite flows have erupted. Right lateral motion has produced extensional pull-apart deformation. Outcrops in the Vega de los Colorados region include Ordovician low-grade metamorphic sedimentary rocks that are unconformably overlain by clastic sediments and mafic lava flows. The age of the sedimentary sequence is constrained by a 40Ar/39Ar age of 26.3 ± 1.6 Ma on an interbedded tuff (Vandervoort et al., 1995). The sequence is affected by folds with NNE-SSW–trending axes that are consistent with subhorizontal WNW-ESE shortening and subvertical extension. One of the basaltic andesitic lava flows has a 40Ar/39Ar plateau age of 4.6 ± 0.2 Ma (Risse et al., 2008; isochron age is 4.9 ± 0.2 Ma). The mafic flows and scoria cone in the Vegas de los Colorados are cut by a fault that strikes 342° and dips 45° northeast. This fault is interpreted as a Tertiary reverse fault reactivated as a normal fault (Marrett et al., 1994). Stop 3-7: Nacimientos—Basaltic Andesites and Their Tectonic Control

25°30′S 14–10 Ma lavas

9–8 Ma lavas

5

4.6 Ma bas

6

9–8 Ma lavas

14–10 Ma lavas 7–4 Ma lavas 10 km Plio-Qbas

Plio-Qbas Pc/Pal

67°30′W

Figure 3.6. Local map of the principal features of region of the Vega de los Colorados showing locations of Stops 3-5 and 3-6. Pc/Pal is Precambrian–lower Paleozoic basement.

Directions: Take dirt road from Vega de los Colorados south in the direction of Nacimientos. Stop is at 25°52′13.9″; 67°26′9.3″; ~3750 m asl. Basaltic andesite lavas with olivine and clinopyroxene phenocrysts at this stop are typical of flows erupted from monogenetic and simple polygenetic cones in the Nacimientos volcanic field of the southern Puna (Figs. 3.1 to 3.3). This flow yielded a 40Ar/39Ar age of 2.8 ± 0.2 Ma (Risse et al., 2008). The source of the flow is a small monogenic center located on a northnortheast–striking fault system. The fault shows right-lateral movement, which produced an extensional pull-apart setting that controlled the location of the volcanic vents. Stop 3-8: View of Basaltic Andesites and Andesites Centers Directions: Continue on dirt road. Stop is at 25°53′03″S; 67°27′045″W; 3735 m asl. This stop provides a general view of the numerous small basaltic to mafic andesitic centers west of the Cerro Galán caldera. These centers were emplaced along north-south– to northnortheast–striking fault systems. The mafic lavas show a temporal change in chemical character from more arc-like (e.g., La/Ta ratios = 36–55) at ca. 6.6 Ma to more intraplate-like (e.g., La/Ta = 20–35) affinities after 3 Ma (Kay et al., 1994a, 1999). This change is interpreted as being related to the loss of a backarcsubducted component in the thicker mantle wedge that resulted from steepening of the subduction zone and delamination of the crust and mantle.

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DAY 4—SOUTHERN PUNA SILICIC CALDERAS The main objective of Day 4 is to observe the Cerro Galán and Cerro Blanco silicic caldera complexes and their regional settings. The centers are shown in the map in Figure 4.1. Stop 4-1: Early Pliocene Cerro Merihuaca Lava Flows, ca. 2 Ma Galán Ignimbrite, and a View of Mafic Cinder Cones and Lava Flows to the West Directions: Go north on Route 43 from Antofagasta de la Sierra. Turn east at the sign for the road to Real Grande, ~6 km north of the village of Antofagasta de la Sierra; turn is near 26°02′30S; 67°23′041″W; 3441 m asl. Stop is at 25°59′07″S; 67°17′27″W; 3984 m asl.

68°30′WSalar

Dacitic lava flows from the Merihuaca volcanic center (4.86 ± 0.19 Ma; Sparks et al., 1985) are well seen at this stop (Figs. 4.2 and 4.3). They are surrounded by the Cerro Galán ignimbrite. The surface of the ignimbrite viewed from this stop exhibits a distinctive wavy form that reflects its eruption on a smooth slope dropping into the Punilla River depression to the west. Numerous late Miocene to Pleistocene basaltic andesitic, basaltic trachyandesitic, and andesitic monogenic and polygenic centers can be seen to the west (Figs. 4.1 and 4.2). Stop 4-2: Wavy Upper Surface of the Galán Ignimbrite Directions: Continue on dirt track; stop is at 25°58′56″S; 67°17′08″W; 4021 m asl.

67°30′W

66°30′W

Aguas Calientes

l l a a

Volcan Antofalla

l a

r

A

n

t

o

f

Agua Escondida

25°30′S

Laguna Los Patos

a

Caldera C° Galan

S

Los Colorados Ig.

Cal

alas

te

26°00′S

Sie

rra

de

Salar de Ratones

Vallecito Ig.

Antofagasta de la Sierra

Salar de Incahuasi

N

LEGEND Lagoon Salars C° Blanco Caldera Complex: