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 <~dit(\d

by Suzanne lahlbur·g Ka)· and

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A. Han1os

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

edited by Suzanne Mahlburg Kay Cornell University Department of Earth and Atmospheric Sciences Snee Hall Ithaca, New York 14853, USA Víctor A. Ramos Universidad de Buenos Aires Laboratorio de Tectónica Andina Consejo Nacional de Investigaciones Científicas y Técnicas (CONICET) Buenos Aires 1428, Argentina

Field Guide 13 3300 Penrose Place, P.O. Box 9140

Boulder, Colorado 80301-9140 USA

2008

Copyright © 2008, The Geological Society of America (GSA). All rights reserved. GSA grants permission to individual scientists to make unlimited photocopies of one or more items from this volume for noncommercial purposes advancing science or education, including classroom use. For permission to make photocopies of any item in this volume for other noncommercial, nonprofit purposes, contact The Geological Society of America. Written permission is required from GSA for all other forms of capture or reproduction of any item in the volume including, but not limited to, all types of electronic or digital scanning or other digital or manual transformation of articles or any portion thereof, such as abstracts, into computer-readable and/or transmittable form for personal or corporate use, either noncommercial or commercial, for-profit or otherwise. Send permission requests to GSA Copyright Permissions, 3300 Penrose Place, P.O. Box 9140, Boulder, Colorado 80301-9140, USA. Copyright is not claimed on any material prepared wholly by government employees within the scope of their employment. Published by The Geological Society of America, Inc. 3300 Penrose Place, P.O. Box 9140, Boulder, Colorado 80301-9140, USA www.geosociety.org Printed in U.S.A. GSA Books Science Editors: Marion E. Bickford and Donald I. Siegel Library of Congress Cataloging-in-Publication Data Field trip guides to the backbone of the Americas in the southern and central Andes : ridge collision, shallow subduction, and plateau uplift / edited by Suzanne Mahlburg Kay, Víctor A. Ramos. p. cm. — (Geological Society of America ; field guide 13) Includes bibliographical references. ISBN 978-0-8137-0013-7 (pbk.) 1. Geology, Structural—Andes Region—Guidebooks. 2. Geology—Andes Region— Guidebooks. I. Kay, Suzanne Mahlburg, 1947–. II. Ramos, Víctor A. QE230.F54 2008 558—dc22 2008017652

Cover: Mount Aconcagua, highest peak in the western hemisphere. Late Miocene volcanic sequences form the top of the mountain.

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Contents

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . v 1. Field trip guide: Ridge-trench collision—The southern Patagonian Cordillera east of the Chile Triple Junction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1 Matthew L. Gorring 2. Field trip guide: Andean Cordillera and backarc of the south-central Andes (~38.5°S to 37°S). . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23 Tomás Zapata, Gonzalo Zamora Valcarce, Andrés Folguera, and Daniel Yagupsky 3. Field trip guide: Frontal and Main Andean Cordilleras near the southern boundary of the Pampean shallow subduction zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57 Víctor A. Ramos 4. Field trip guide: Evolution of the Pampean flat slab region over the shallowly subducting Nazca plate . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 77 Víctor A. Ramos 5. Field trip guide: Neogene evolution of the central Andean Puna plateau and southern Central Volcanic Zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 117 Suzanne Mahlburg Kay, Beatriz Coira, and Constantino Mpodozis

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The Geological Society of America Field Guide 13 2008

Introduction Suzanne Mahlburg Kay* Department of Earth and Atmospheric Sciences, Snee Hall, Cornell University, Ithaca, New York 14853 USA Víctor A. Ramos* Laboratorio de Tectónica Andina, Universidad de Buenos Aires, Consejo Nacional de Investigaciones Científicas y Técnicas (CONICET), Buenos Aires 1428, Argentina The field guides in chapters 1–5 of this volume were originally written to accompany the field trips run in conjunction with the Backbone of the Americas Meeting that took place 3–7 April 2006, in Mendoza, Argentina. The conference was sponsored by the Geological Society of America (GSA) and the Asociación Geológica Argentina (AGA) in collaboration with the Chilean Geologic Society. The meeting was organized around three important processes affecting the western margin and cordilleras of the Americas—ridge collision, shallowing and steepening subduction zones, and plateau and orogenic uplift. Some 400 participants from 20 countries along the western margin of the Americas and elsewhere attended the meeting. A report of the meeting can be found in the January 2007 issue of GSA Today (Kay and Ramos, 2007). The five field guides from the pre-meeting, intra-meeting, and post-meeting field trips in this volume accompany the meeting themes. General trip locations are shown in the map and topographic image in Figure 1. The trip in chapter 1 to southern Patagonia highlights the ridge-trench collision theme; the trips in chapters 2–4 examine issues related to shallowing subduction zones in the south-central Andes; and the trip in chapter 5 highlights uplift of the Puna plateau in the context of a steepening subduction zone and lithospheric delamination. Chapter 1, “Ridge-trench collision—The southern Patagonian Cordillera east of the Chilean Triple Junction,” is the guide for premeeting field trip 401 to Patagonia, which was organized to highlight the ridge collision theme. This five-day field guide by Matthew Gorring (Montclair State University) features the southern Patagonian Cordillera, south of the present Chile Triple Junction (46.5°S), where distinctive backarc deformational and magmatic features reflect the northward-propagation of slab windows related to the ridge-trench collision event that began at ca. 14 Ma. Among the features to be observed are (1) an abrupt northward increase in topography in the Cordillera, (2) young thin- and thick-skinned deformation in the Patagonian fold-thrust belt, (3) mid-Miocene to Pliocene plutons associated with the shutoff of the arc south of the slab window, (4) adakitic volcanic rocks interpreted as slab melts associated with partial melting of the hot trailing edge of the Nazca plate as the slab window opened, and (5) extensive oceanic island basalt–like plateau basalts formed from dynamic asthenospheric flow as the slab window opened under the backarc. Chapter 2, “Andean Cordillera and backarc of the south-central Andes (~38.5° to 37°S),” is the guide for pre-meeting field trip 402 to the Neuquén Andes in Argentina, which was organized to highlight a region of the margin where both contractional and extensional deformation have occurred. This five-day field guide by Tomás Zapata of Repsol-YPF, Gonzalo Zamora (Repsol-YPF), Andrés Folguera (University of Buenos *E-mails: [email protected]; [email protected]

Kay, S.M., and Ramos, V.A., 2008, Introduction, in Kay, S.M., and Ramos, V.A., eds., Field trip guides to the Backbone of the Americas in the southern and central Andes: Ridge collision, shallow subduction, and plateau uplift: Geological Society of America Field Guide 13, p. v–vii, doi: 10.1130/2008.0013(00). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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Kay and Ramos 80° W

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plateau Nazca Plate

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idge zca R

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Figure 1. Map and topographic relief images of South America show the locations of the five field trips. The numbers 1–5 refer to the chapters in the field guide.

Aires), and Daniel Yagupsky (University of Buenos Aires) features the Cretaceous to Recent deformational history of the Andes adjacent to the Neuquén Mesozoic sedimentary basin. The region has experienced multiple deformational events in response to changes in the configuration of the convergent plate boundary. Among the features to be observed are (1) the Mesozoic statigraphic sequence of the western Neuquén basin, (2) the uppermost Cretaceous Agrio fold-and-thrust belt whose origin is linked to moderate shallowing of the subduction zone by the authors, (3) the late Oligocene–early Miocene Cura Mallín extensional basin related by the authors to steepening of the subducting plate, (4) contractional structures related to the inversion of the Cura Mallín basin and the formation of the Guañacos fold and thrust belt, (5) mafic volcanic rocks and extensional structures in the Loncopué trough related to recent steepening of the subduction zone, and (6) the spectacular young backarc Cerro Troman volcano. Chapter 3, “Frontal and Main Andean Cordilleras near the southern boundary of the Pampean shallow subduction zone,” is the guide for field trip 403, which was the intra-meeting field trip to the high central Andes for all participants. This one-day field guide by Víctor Ramos (University of Buenos Aires) is for a transect above the currently amagmatic southern hinge of the Pampean flat slab (commonly called the Chilean flat slab). The trip provides a view of the Andean deformation front in the Precordillera and the main features of the Frontal and Main Cordilleras in Argentina. Among features to be observed are (1) late Paleozoic sedimentary and magmatic rocks in the Frontal Cordillera, (2) Triassic volcanic and plutonic sequences, (3) the sedimentary sequences and inverted normal structures of the Triassic Cuyo rift, (4) Mesozoic to Miocene magmatic arc and sedimentary basin rocks in the Main Cordillera, (5) Miocene foreland basin deposits, (6) the thin-skinned Aconcagua fold and thrust belt that deforms Jurassic to Miocene sequences, and (7) the late Miocene Cerro Aconcagua center (6967 m)—the highest peak in the Western and Southern hemispheres (the top of the Backbone). Chapter 4, “Evolution of the Pampean flat-slab region over the shallowly subducting Nazca plate,” is the post-meeting guide for field trip 404, which was organized to highlight the shallowing subduction zone

Introduction

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theme. This five-day field guide by Víctor Ramos (University of Buenos Aires) provides a view of the Paleozoic to Recent history of the region over the Pampean flat-slab region (also known as the Chilean flat slab). The trip features the Andes between 31° and 32°S latitude in a currently amagmatic segment characterized by flat-slab subduction. Among features highlighted are (1) neotectonic activity at the boundary between the basement uplifts of the westernmost Sierras Pampeanas and the Precordillera belt, (2) Precordillera belt Paleozoic stratigraphy and evidence for Ordovician accretion of the Cuyania (greater Precordillera) terrane and Devonian accretion of the Chilenia terrane, (3) Neogene structure and sedimentary and volcanic rocks of the Precordillera fold and thrust belt, (4) western Precordillera late Paleozoic Gondwana sequences and Triassic rift sediments and structures, and (5) a transect from the Frontal Cordilleras to the Coastal Cordillera in Chile that complements the guide in Chapter 3. Chapter 5, “Neogene evolution of the central Andean Puna plateau and the southern Central Volcanic Zone,” is the guide for post-meeting field trip 405, which was organized to highlight the plateau uplift theme as well as a steepening subduction zone and crustal and lithospheric delamination. This seven-day field guide by Suzanne Kay (Cornell University), Beatriz Coira (Universidad Nacional de Jujuy, Argentina; Days 1–5) and Constantino Mpodozis (Antofagasta Minerals, Chile; Days 6–7) provides a view of plateau evolution in the context of the southern central Andean Puna–Altiplano plateau in Argentina between 23° and 27.5°S latitude and the southernmost Central Andean Volcanic Zone arc in Chile between 26.5° and 27.5°S latitude. Differences in the evolution and uplift of the northern and southern Puna are highlighted on the first five days. Featured are (1) the giant Miocene northern Puna Coranzulí ignimbrite and the Pliocene to Pleistocene southern Puna Cerro Galan and Cerro Blanco ignimbrites, (2) the central Puna Pleistocene shoshonitic lavas and the southern Puna latest Miocene to Pleistocene intraplate and calc-alkaline lavas, (3) latest Miocene to Recent southern Puna normal and strike-slip faults, (4) Tertiary sedimentary sequences and post-middle Miocene internally drained salar basins, (5) Ordovician to Cretaceous basement and deformational styles, and (6) the structure of the southeastern margin of the plateau. The last two days highlight the late Miocene to Pliocene displacement of the arc front from the Maricunga Belt to the southern Central Volcanic Zone. Featured are (1) the Miocene Maricunga Belt arc front centers on the western margin of the plateau, (2) early Miocene mafic rocks and Miocene-Pliocene volcanic rocks between the Maricunga Belt and Central Volcanic Zone, (3) evidence for early to middle Miocene backarc contractional deformation, and (4) the dramatic Central Volcanic Zone volcanic centers near the international border including Ojos del Salado and Tres Cruses. ACKNOWLEDGMENTS We would like to thank the reviewers and trip participants who made significant contributions to improving the presentation and content of these guides. Special thanks go to Dwight Bradley (U.S. Geological Survey), Eric Erslev (Colorado State University), Alan Glazner (University of North Carolina), Teresa Jordan (Cornell University), Robert Kay (Cornell University), and Constantino Mpodozis (Antofagasta Minerals, Chile). Repsol-YPF is acknowledged for offsetting some of the field trip costs. REFERENCE CITED Kay, S.M., and Ramos, V.A., 2007, Backbone of the Americas: From Patagonia to Alaska—A super rock star event: GSA Today, v. 17, no. 1, p. 30–31. MANUSCRIPT ACCEPTED BY THE SOCIETY 10 JANUARY 2008

Printed in the USA

The Geological Society of America Field Guide 13 2008

Field trip guide: Ridge-trench collision—The southern Patagonian Cordillera east of the Chile Triple Junction Matthew L. Gorring* Montclair State University, Montclair, New Jersey 07043, USA

ABSTRACT The southern Patagonian Cordillera south of the present location of the Chile Triple Junction (46.5°S) preserves distinctive deformational and backarc magmatic features that are a consequence of a series of northward-propagating ridge collision events that started at ca. 14 Ma. An abrupt increase of ~2000 m of topographic elevation and the exhumation and uplift of mid-Miocene to Pliocene plutons within the cordillera south of the Chile Triple Junction is accomplished by horizontal compressive deformation (both thin- and thick-skinned) within the Patagonian fold-thrust belt. Ridge-trench collisions have formed asthenospheric slab windows beneath the southern Patagonian Cordillera. Backarc magmatism associated with slab window formation includes a distinctive suite of adakites and extensive outpourings of oceanic island basalt (OIB)-like plateau basalts. The adakites formed from the partial melting of the young, hot trailing edge of the Nazca plate that preceded slab window opening, whereas the OIB-like plateau basalts formed from dynamic asthenospheric flow as the slab windows opened up beneath the backarc. Keywords: Patagonia, slab window, Chile Triple Junction, ridge-trench collision, southern Andes. INTRODUCTION

Guivel et al., 1999), anomalous forearc felsic and mid-oceanic ridge basalt (MORB)-like magmatism (Mpodozis et al., 1985; Forsythe et al., 1986; Lagabrielle et al., 1994; Le Moigne et al., 1996), production of a gap in the Quaternary volcanic arc (Stern et al., 1990), rapid uplift and reactivation of deformation in the fold-and-thrust belt (Ramos, 1989; Coutand et al., 1999; Kraemer et al., 2002), and adakitic and OIB-like basaltic magmatism in the backarc (Gorring et al., 1997; Ramos et al., 2004). All of these well-documented effects, coupled with the relatively simple and well-constrained plate convergence geometry, make southern Patagonia one of the best places on Earth to investigate geodynamic processes associated with ridge collision.

The purpose of this field trip is to examine the geologic features associated with the process of ridge-trench collision along a partial transect across the southern Patagonian Cordillera near the latitude of the present location of the Chile Triple Junction at 46.5°S latitude (Fig. 1). The effects of the northward-propagating ridge collision events starting at ca. 14 Ma at 54°S latitude are well documented by ophiolite emplacement and forearc subduction erosion (Bourgois et al., 1996; Mpodozis et al., 1985; *[email protected]

Gorring, M.L., 2008, Field trip guide: Ridge-trench collision—The southern Patagonian Cordillera east of the Chile Triple Junction, in Kay, S.M., and Ramos, V.A., eds., Field trip guides to the Backbone of the Americas in the southern and central Andes: Ridge collision, shallow subduction, and plateau uplift: Geological Society of America Field Guide 13, p. 1–22, doi: 10.1130/2008.0013(01). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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Figure 1. Tectonic setting of southern South America showing the distribution of Neogene plateau lavas, adakites, the Patagonian fold-and-thrust belt, and the Southern (SVZ) and Austral Volcanic Zones (AVZ). Ridge collision times are shown in bold numbers (Cande and Leslie, 1986).

However, due to the geography, difficult access, and long distances, this trip will concentrate on features that are most easily accessible from the Argentine foreland basin and backarc side, into the hinterland of the cordillera in Chile, but not as far as the forearc region in the Taitao Peninsula area. The main highlights of this trip will include spectacular, “whole mountain-side” exposures of the hinterland and foreland portions of the Patagonian fold-thrust belt and examples of the Late Miocene to Pleistocene adakitic and basaltic magmatism that occurred in the backarc. Along the way, we will also observe other classic geologic features of the southern Patagonian Andes including the Miocene Fitz Roy pluton, Jurassic rhyolites, late Paleozoic low-grade metamorphic basement, Cretaceous granitoids of Patagonian Batholith, well-exposed Cretaceous and Tertiary sedimentary sections, and Pleistocene moraine complexes and geomorphology. Users of this guide should keep in mind that this area presents particular challenges in terms of logistics. Air access, vehicle rental, and availability of modern amenities are generally limited to the three major towns and cities in the region (El Calafate, Comodoro Rivadavia in Argentina, and Coyhaique in Chile). Outside of these major towns, fuel is generally available in the small towns and outposts, but distances are deceptively great and the region desolate. Therefore, one must stop at every available gas station to fuel up and carefully plan for adequate supplies of food and water. Lodging is generally not an issue, but don’t expect modern facilities outside of the major towns. The majority of the roads are gravel but are well maintained. Gravel eventu-

ally will take its toll on the vehicle, thus one must have at least two “real” spare tires and an array of essential tools (heavy-duty jack, assorted wrenches, shovel, etc.) to deal with problems on the road. Groceries and lodging are generally quite limited outside the major towns. THE SOUTHERN PATAGONIAN CORDILLERA The Andean Cordillera lines the entire 7000-km-long western margin of South America where subduction of oceanic plates has occurred beneath the continent. They are the result of crustal thickening due to horizontal shortening and magmatic addition since the Late Cretaceous. The southern Patagonian Cordillera is defined here as the segment of the Andes, south of the latitude of the Chile Triple Junction at 46.5°S and north of 53°S where the more E-W–trending Fueguian Cordillera approximately begins. From east to west, the southern Patagonian Cordillera can be divided into four distinctive zones (e.g., Kraemer, 1993; Diraison et al., 2000; Ramos, 2005). 1. The Magallanes (or Austral) foreland basin, containing relatively undeformed early Cretaceous to Tertiary sediments reaching a maximum thickness of 8000 m. 2. The Patagonian fold-and-thrust belt, where late Paleozoic sedimentary and metamorphic rocks, Jurassic volcanics, and early Cretaceous to Tertiary sediments form belts of thickand thin-skinned deformation.

Ridge-Trench Collision—The Southern Patagonian Cordillera 3. The central main cordillera, where late Paleozoic sedimentary and metamorphic rocks, Jurassic volcanics, and Mesozoic to Miocene calc-alkaline granitoids of the Patagonian Batholith are characterized mainly by thick-skinned, high-angle reverse and thrust deformation. 4. The western coastal belt, where a highly deformed, accretionary prism complex of late Paleozoic metasedimentary rocks are intruded by granitoids of the Patagonian Batholith. Late Paleozoic Paleozoic basement rocks are widely exposed in the southern Patagonian Cordillera and record evidence for a major episode of crustal shortening along the margin of the supercontinent Gondwana. They consist of two different units: (1) lowgrade metasedimentary rocks of the Rio Lacteo Formation and (2) the thick sedimentary sequence of the Bahia La Lancha Formation. Rocks of the Rio Lacteo Formation are found to the north, between 45° and 48°S, and are dominated by quartzites, mica schists, and phyllites with minor greenstones and marbles (Leanza, 1972; Hervé, 1988; Hervé et al., 1998; Bell and Suárez, 2000). The Bahia La Lancha Formation is a typical flysch sequence of graywacke, quartz sandstones, and shale up to 2000 m thick that is exposed to the south between 48° and 51°S (Riccardi, 1971). Time constraints for the Rio Lacteo Formation are derived from K-Ar ages of intrusive tonalitic rocks of Late Carboniferous–Early Permian at various locations (Halpern, 1973; Niemeyer et al., 1984; Ramos, 1989; Bell and Suárez, 2000). Bahia La Lancha rocks contain fossils of Late Devonian–Carboniferous age (Riccardi, 1971). Main deformation and metamorphism of both units is thought to be related to the docking of various allochthonous terranes that collided with the active Pacific margin of Gondwana in the latest Paleozoic (Gondwanide Orogeny) (Ramos, 1988). Triassic to Middle Jurassic Following the period of terrane accretion in the late Paleozoic, the southwestern margin of Gondwana appears to have experienced widespread extension and crustal thinning that preceded the eventual breakup of the supercontinent (Dalziel, 1981). From the late Triassic to mid-Jurassic, a system of NNWoriented rift basins were formed, and vast quantities of silicic volcanics (mainly rhyolite and associated volcaniclastics) covered more than 106 km2 of Patagonia (Gust et al., 1985; Pankhurst et al., 1998). This important suite of volcanics in southern South America is known as the Chon Aike province (Kay et al., 1989; Pankhurst et al., 1998). Locally these rocks have various formation names such as Tobifera Formation (Thomas, 1949), El Quemado Complex (Riccardi, 1971), and Ibáñez Formation (Niemeyer et al., 1974). Chon Aike volcanics were deposited unconformably over the late Paleozoic basement and locally attain thicknesses of >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.

6

Gorring

• 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 ).

10

Gorring

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

11

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 <0.1 Ma), less voluminous, alkaline (43%–49% SiO2; 5%–8% Na2O + K2O) post-plateau sequence (Fig. 7). The main-plateau lavas form large, elevated

TABLE 1. GEOCHEMICAL AND AGE DATA FOR PATAGONIAN SLAB WINDOW VOLCANICS Cerro Pampa Puesto Nuevo Chaltén Post-plateau PNA1 48°56.0′ 72°12.5′

FVR 49°25.5′ 72°59.5′

LC-7 49°11.4′ 71°20.7′

Main-plateau

Sample Latitude (°S) Longitude (°W)

RB5 47°54.5′ 71°25.4′

RB7 47°54.5′ 71°25.4′

RB8 47°54.5′ 71°25.4′

LC-21 48°18.3′ 70°58.2′

SiO2 (wt%) TiO2

62.55 0.62

67.88 0.51

63.22 0.84

65.67 0.68

64.97 0.66

47.89 2.96

53.22 1.60 15.96

Al2O3

17.27

16.26

16.53

16.56

15.21

15.11

FeO

3.13

2.24

3.78

3.02

3.32

11.40

8.70

MnO

0.09

0.11

0.11

0.10

0.06

0.17

0.13

MgO

3.62

2.48

2.92

2.92

3.42

7.22

6.21

CaO

7.12

4.50

6.72

4.40

4.51

7.89

8.14

Na2O

4.66

4.56

4.31

4.25

3.66

4.07

4.02

K2O

1.22

1.70

1.88

2.70

3.70

1.77

0.48

P2O5 Total

N.D. 100.3

N.D. 100.2

N.D. 100.3

0.25 100.6

0.24 99.8

1.02 99.5

0.22 98.7

La (ppm)

27

21

40

27

30

49

9

Ce

61

45

97

64

64

93

19

Nd

30

20

52

29

27

43

13

Sm

4.4

3.2

7.6

4.9

5.0

8.8

3.8

Eu

1.16

0.77

1.93

1.16

1.19

2.63

1.34

Tb

0.34

0.27

0.63

0.38

0.43

1.06

0.66

Yb

0.72

0.69

1.20

0.96

0.99

1.72

1.50

Lu

0.09

0.08

0.16

0.12

0.13

0.21

0.23

Sr

1886

1334

2294

1439

1369

925

452

Ba

306

391

356

320

381

523

135

Cs

0.30

0.30

0.60

3.20

1.00

0.60

Rb

N.D.

N.D.

N.D.

N.D.

N.D.

N.D.

7.3

U

1.00

2.60

1.80

2.00

3.10

1.75

0.54

6.4

1.1

14.9

0.20

Th

5.0

6.7

7.5

7.0

Pb

N.D.

N.D.

N.D.

N.D.

N.D.

N.D.

Y

N.D.

N.D.

N.D.

N.D.

N.D.

N.D.

20

Zr

N.D.

N.D.

N.D.

N.D.

N.D.

N.D.

108

Hf

3.2

3.2

5.0

5.4

5.7

5.8

Nb

N.D.

N.D.

N.D.

N.D.

N.D.

N.D.

Ta

0.70

1.80

0.70

Sc

8

6

0.50

0.80

11

9

9

4.39 17

1.5

2.4 13 0.84 18

Cr

97

85

81

100

102

159

218

Ni

76

47

43

54

68

116

133

Co 87 86 Sr/ Sr

15 0.70290

15 0.70285

14 0.70321

13 0.70330

εNd Ar/Ar age (Ma)

+5.7

+6.9

+4.9

~+5

13.1 ± 0.6

14.5 ± 0.3

N.D.

9 0.70309 +5.5 N.D.

11.4 ± 0.6

48 N.D.

40 0.70358

N.D.

N.D.

<6 (?)

ca. 12

Note: Adakite data from Kay et al. (1993) and Ramos et al. (2004); slab window basalt data from Gorring and Kay (2001). N.D.—not determined.

Figure 7. (A) Map showing the distribution of Neogene main-plateau and post-plateau lavas, and the location of the Cerro Pampa adakite. AVZ— Austral Volcanic Zone; SSVZ—southern Southern Volcanic Zone. (B) Plot of slab window plateau lavas ages projected onto the SW-NE transect in (A) showing the northeastward younging trend of the main- and post-plateau lavas. Arrow shows predicted age progression for a mantle hotspot track based on a 2.5 cm/yr absolute plate motion vector for South America. Gray zone shows the predicted time when the trailing edge of the Nazca plate passed beneath the backarc. Both figures modified from Gorring et al. (1997).

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Gorring

plateaus of the Mesetas de la Muerte, Belgrano, del Lago Buenos Aires, Central, and the smaller mesetas in the northeast region. The main-plateau lavas have a maximum total thickness of 100–200 m, and their total eruptive volume is estimated at 1000–2000 km3. The post-plateau lava sequence includes small scoria cones, lava flows, and pyroclastic deposits capping the main-plateau sequence with a total eruptive volume estimated at 100 km3. Both main- and post-plateau lavas have strong OIB-like geochemical signatures (La/Ta < 20; Ba/La < 20; 87 Sr/86Sr = 0.7035–0.7049; 143Nd/144Nd = 0.51290–0.51261) (Table 1) and, in the backarc segment opposite the 12 Ma ridge collision, systematically young to the northeast (Fig. 7B). A slab window tectonic model has been developed by Gorring et al. (1997) to explain the sequence of magmatic events occurring along a SW-NE transect opposite the Chile Ridge segment that collided at ca. 12 Ma as shown in Figure 2. Normal arc volcanism, similar to that in the modern Andean Southern Volcanic Zone, is assumed to have occurred in the modern arc gap prior to ridge collision in the early Miocene. By 12 Ma, the ridge had collided, causing volcanism in the former arc region to cease and shift into the backarc. Partial melting of the young, hot, Nazca Plate produced the southern Patagonian adakites. Contemporaneous eruptions of minor volumes of mafic backarc lavas are best related to melting in the supraslab mantle wedge that preceded the opening of the slab window. From 10 to 2 Ma, mafic lavas with strong OIB-like chemical signatures derived from the subslab asthenospheric mantle erupted across the backarc above a developing slab window. Main-plateau lavas represent relatively large-degree partial melts (10%–15%) associated with strong asthenospheric flow around the trailing Nazca plate edge, whereas post-plateau lavas are small-degree melts (1%–4%) generated by weak asthenospheric flow through a wide slab window. A simple one-stage melting model of a pristine OIB-like subslab asthenospheric source can explain most post-plateau lavas. Main-plateau lavas from the western backarc are also dominated by the OIBlike subslab asthenospheric mantle, but clearly have evidence for arc and crustal components (relatively low Nb/U and Ce/Pb and high Sr/La) derived from interaction with slab-melt and fluidcontaminated mantle wedge and continental lithosphere. Finally, backarc magmatism ends and the locus magmatic activity shifts westward to form the arc volcanism in the Austral Volcanic Zone above the leading edge of the Antarctic plate (Stern and Kilian, 1996). Farther north, opposite the ridge segment that collided at 6 Ma, a similar dynamic slab window model can explain abundant Plio-Pleistocene (5–0.2 Ma), main- and post-plateau lavas that occur in the Meseta del Lago Buenos Aires region. ROAD LOG Day 1—El Calafate to El Chaltén We will drive east out of Calafate on Provincial Route 11. On the south side of the valley, one can observe (from a distance) the beautifully exposed (glacially carved), upper Cretaceous and

Tertiary foreland basin sedimentary section. We then turn north on the famous “Ruta Cuarenta” or National Route 40 and soon pass over the Rio Santa Cruz and cross over the Rio La Leona. After ~70 km, turn left onto Provincial Route 23 and travel west toward the cordillera along the north side of Lago Viedma. We will make several stops along this road to examine the large-scale structure of the Patagonian fold-and-thrust belt as we drive west into the town of El Chaltén. On the north side of the valley, Late Cretaceous and Tertiary sediments are unconformably overlain by flat-lying basalts that form high plateaus (Mesetas Basaltica, Chica, and del Viento). These plateau basalts have early Pliocene K/Ar whole-rock ages of ca. 3.5 Ma (Mercer, 1976). In this area, we are located near the eastern edge of the deformation front in the foreland basin where the Cretaceous sediments are only mildly deformed and fold wavelengths are large (km scale) and amplitudes are small. Stop 1-1: Patagonian Fold-Thrust Belt Once west of the Rio Cangrejo, deformation intensity noticeably increases as we pass a section of the late Cretaceous La Anita and Pari Aike Formations that are moderately east dipping beneath the western edge of the Meseta Chica. The first stop(s) will be at or near the bridge over the Rio Blanco, where we can observe (from a distance) structures associated with surface exposure of the easternmost, thin-skinned thrusts faults (both east and west verging) (Figs. 8 and 9). Folds in the marine sandstones of the Cerro Toro Formation beneath the Bardas de Kaiken Aike are of chevron, kink, and concentric styles, typical of flexural slip in rhythmically layered, turbidite sequences. Stop 1-2: Patagonian Fold-Thrust Belt The second major stop will be on the east side of the bridge across the Rio de las Vueltas. Along the gorge walls, there are very nice exposures of the early Cretaceous black shales of the Rio Mayer Formation cut by several small basaltic(?) dikes and capped by glacial deposits. The black shales of the Rio Mayer Formation form the basal décollement for the thin-skinned part of the Patagonian fold-and-thrust belt. On the south-facing flanks of Cerro Faldeo there are spectacular folds (and thrusts) within the Rio Mayer Formation. Stop 1-3: Patagonian Fold-Thrust Belt The third stop will be a few kilometers to the west at a formal scenic viewpoint just within the national park boundary. This vantage point, on a clear day, affords a spectacular view of the granitic spires of Cerro Fitz Roy (3405 m) and Cerro Torre (3102 m) and an overview of the thick-skinned portion of the Patagonian fold-and-thrust belt. At this location, we are only a few kilometers east of the first of several east-verging, steeply dipping, thick-skinned thrusts that uplift and bring to the surface late Paleozoic marine metasediments (low-grade Bahia La Lancha Formation, which forms the basal décollement), Jurassic volcanics of the El Quemado Complex, early Cretaceous shales of the Rio Mayer Formation, and the Miocene

Ridge-Trench Collision—The Southern Patagonian Cordillera

15

Co Faldeo Bardas de Kaiken Aike

Meseta Basaltica

Figure 8. Geologic and structural map of the northern shore of Lago Viedma, with location of the structural section in Figure 9. Figure from Coutand et al. (1999).

Figure 9. Structural cross section partly controlled by seismic data. Black bars indicate approximate location of seismic lines. No vertical exaggeration. Figure from Coutand et al. (1999).

Cerro Fitz Roy pluton (Figs. 8 and 9). Depending on weather conditions and time, several stops will be made near the town of El Chaltén or in the Lago Desierto region to observe the general geology of the area. Overnight is in El Chaltén at the Hotel Fitz Roy. Day 2—El Chaltén to Estancia La Angostura Stop 2-1: Cerro Fitz Roy Pluton Weather permitting, we will take a short trek (~6 km roundtrip; ~200 m elevation change) up to the first good, relatively close viewpoint of the Cerro Fitz Roy pluton. On the way, we will be walking over and around outcrops of the Jurassic El Quemado

Formation, which is a thick series of rhyolitic volcanics. The Miocene Cerro Fitz Roy pluton is a composite calc-alkaline pluton made up of mostly coarse-grained quartz monzonite and diorite and has an imprecise K/Ar age of 18 ± 3 Ma (Nullo et al., 1978). The pluton intrudes the Jurassic El Quemado Complex and the late Paleozoic Bahia La Lancha Formation and generally has steep, sharp contacts that can be seen in the high col between Aguja Saint Exupery and Mojon Rojo and the high col between southeast flank of Cerro Torre and the Cordon Adela. To the south of Cerro Fitz Roy, Coutand et al. (1999) identified an associated granitic sheet that has been thrust over intensely deformed Rio Mayer shales indicating that the Cerro Fitz Roy pluton is a syntectonic intrusion.

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Stop 2-2: Chaltén Adakite After the trek, we will drive east out of El Chaltén back along Provincial Route 23. The second stop of the day will be the Chaltén adakite (Ramos et al., 2004), located ~10 km east of El Chaltén, just inside the national park boundary. The outcrop forms a low hill on the north side of Route 23 that is partially covered by glacial till. The Chaltén adakite is a coarse-grained, porphyritic dacite with abundant acicular to prismatic hornblende phenocrysts (2–3 cm) in a matrix of fine-grained plagioclase and clinopyroxene. The Chaltén adakite has all the classic geochemical characteristics of a slab-melt adakite (high SiO2, Sr, Cr, Ni, steep REE patterns, and MORB-like 87Sr/86Sr and 143Nd/144Nd ratios) (Table 1, sample FVR; Fig. 10). Ramos et al. (2004), obtained 40Ar/39Ar laser ablation ages on hornblende phenocrysts that yielded a step-heated plateau age of 14.5 ± 0.29 Ma. This age predates ridge collision by ca. 2.5 Ma and is consistent with partial melting of the young portion of the Nazca plate that would have been subducting beneath backarc at this time (see model in Fig. 2).

Stop 2-3: Slab Window Plateau Basalts Continue eastward on Route 23 and then turn north on National Route 40 toward the small town of Tres Lagos. Continue north on Route 40 from the gas station at Tres Lagos. About 20 km north of Tres Lagos, the road begins to enter an increasingly narrower valley with basalt plateaus on both sides. These are late Miocene to Pliocene alkali basalts of the post-plateau sequence of the Meseta La Siberia and are dated at ca. 5–7 Ma (Gorring et al., 1997). The road gradually climbs in elevation and then comes over a pass a few kilometers north of the Estancia La Lucia. The third major stop will be along National Route 40 ~50 km north of Tres Lagos and 2–3 km north of Estancia La Lucia. At this locality, several flows from a young, monogenetic cinder cone complex (Cerro Cordon) located 1–2 km to the east are exposed in several nice roadcuts. Stern et al. (1990) reported small (<5 cm diameter) spinel lherzolite xenoliths in some of the lavas from this locality. These lavas belong to the latest Miocene to Pleistocene

Figure 10. Geochemical plots showing the strong oceanic island basalt (OIB)-like characteristics of Neogene Patagonian main-plateau (filled symbols) and post-plateau (open symbols) slab window lavas. Data and fields from Kay et al. (1993), Gorring and Kay (2001), and Ramos et al. (2004). BSE—bulk silicate earth; HIMU—high U/Pb mantle; MORB—mid-oceanic ridge basalt; NMORB—normal mid-oceanic ridge basalt; OIB—oceanic island basalt; SSVZ—southern Southern Volcanic Zone.

Ridge-Trench Collision—The Southern Patagonian Cordillera post-plateau sequence of Gorring et al. (1997). The lavas are highly alkaline basalts with strong OIB-like characteristics, typical of southern Patagonian slab window plateau lavas (Table 1, sample LC-7; Fig. 10). This locality has not been dated, but other post-plateau lavas from the area clearly postdate the timing of ridge collision by at least 5 Ma and erupted when the slab window beneath this part of the backarc was fully developed (Fig. 2). Stop 2-4: Lago Cardiel Overview Continue north on Route 40 for another ~50 km and stop along the road at overlook of Lago Cardiel. From this vantage point looking west, one can see the inverted topographic relations of the older, higher, mid-Miocene main-plateau lavas (10–12 Ma) of the Mesetas la Siberia and de la Muerte and the early Pliocene (4–5 Ma) post-plateau lavas that cap the low elevation surfaces that form peninsulas on the north and south sides of the lake. Continue north on Route 40 for another ~80 km. Overnight is at the Estancia La Angostura. Day 3—Estancia La Angostura to Lago Posadas Stop 3-1: Rio Chico Overview; Main Plateau Lavas Continue north on Route 40 in the valley of the Rio Chico. The first stop will be above the bridge over the Rio Chico. Here we can obtain good views of mid-Miocene main-plateau lavas to the west (Meseta de la Muerte) and to the north and east (Meseta Central). Pliocene post-plateau lavas can be observed at distinctly lower elevation on the east side of the Rio Chico valley. Beautiful examples of alluvial terrace levels are exposed from the active downcutting of the river. The terraces are composed of the famous “Patagonian gravel,” which forms a more or less continuous sheet (on average a few meters thick, but up to 50–100 m thick in places) across Patagonia. Originally noted and interpreted (more or less correctly) by Charles Darwin, the gravels represent the huge quantities of glacial outwash shed off of the southern Patagonian Cordillera during peak PlioPleistocene glaciations. Stop 3-2: Slab Window Plateau Basalts Continue northwestward on Route 40 for ~35 km. The second stop will be at the road maintenance facility at Tamel Aike. Nice exposures on the north side of the road, behind the buildings, allow us to examine a representative example of the mid-Miocene main-plateau basalts. These lavas are tholeiitic basalts, with flat REE patterns (low La/Yb ratios) and relatively low concentrations of incompatible elements. Although dominantly OIB-like, they have geochemical evidence for contamination with both arc and crustal components (Table 1, sample LC-21; Fig. 10). This particular locality has not been dated, but main-plateau lavas in the immediate vicinity have 40Ar/39Ar total fusion ages of ca. 12 Ma (Gorring et al., 1997) and erupted at about the same time as the ridge collision, and thus, technically are not “slab window lavas” senso stricto.

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Stop 3-3: Glacial Geomorphology Continue north on Route 40 for ~120 km, until we reach the small hamlet of Bajo Caracoles at the head of the glacial valley of Lago Puyrredón. This desolate place sits on top of a moraine that affords spectacular, long distance views of the cordillera crowned by Cerro San Lorenzo (3706 m) and the high, Miocene basalt-capped plateaus to the north (Meseta del Lago Buenos Aires) and south (Meseta Belgrano). Follow Provincial Road 39 west, toward the town of Lago Posadas (also known as Hipólito Yrigoyen). We will cross several prominent moraine complexes with intervening outwash plains. The Lagos Puyreddon and Buenos Aires valleys (the next one to the north) contain a complete record of at least 16 glacial advances since the Greatest Patagonian Glaciation that occurred at ca. 1.1 Ma (Singer et al., 2004). These are some of the best preserved moraine complexes on Earth, let alone in the Southern Hemisphere. Our third stop of the day will be ~35 km west of Bajo Caracoles on Route 39 at the prominent roadcut on the crest of a moraine complex in order to examine its internal structure. Stop 3-4: Patagonian Fold-Thrust Belt As we approach the town of Lago Posadas, we will again see spectacular, glacially carved, continuous, mountainside exposure of the eastern deformation front of the Patagonian fold-and-thrust belt. At the fourth stop, just southeast of the town of Lago Posadas, a complete Cretaceous to early Miocene sedimentary section is exposed on the north flank of the Meseta Belgrano. The entire section is gently tilted to the east into a frontal monocline and marks the eastern limit of deformation (see cross section A–A′ on Fig. 4; section D–D′ on Fig. 11). It is capped by undeformed, mid-Miocene, main-plateau lavas that have been dated ca. 10 Ma (Gorring et al., 1997). From this location, the main peak of Cerro San Lorenzo (3706 m) is also visible, if weather permits. Cerro San Lorenzo is the second highest peak in the Patagonian Cordillera. It is an oval-shaped (130 km2), calc-alkaline granitoid pluton. Recent K/Ar and 40 Ar/39Ar radiometric age dates on various rock types within the pluton have yielded latest Miocene ages ranging from 6.2 ± 0.2 Ma to 6.6 ± 0.5 Ma (Welkner, 2000). Depending on weather conditions and time, additional stops will be made west of the Lago Posadas, in particular a glacially scoured outcrop of Jurassic volcanics, a Younger Dryas(?) moraine that cuts across Lago Puyrredón, and an overlook of the “Garganta del Diablo” on the Rio Oro. Overnight is in Lago Posadas. Day 4—Lago Posadas to Puerto Bertrand via Los Antiguos and Chile Chico Drive east out of Lago Posadas and then north onto the secondary road heading north toward Paso Roballos. This road traverses across the Sierra Colorado, which is a block of the Jurassic rhyolitic volcanics (mostly ignimbrites) of the El Quemado Complex uplifted on east-verging thrust faults (section C–C′ on Fig. 11).

Figure 11. Geologic cross sections of the Patagonian fold-and-thrust belt near Lago Posadas at ~47°S. Figure from Giacosa et al. (1999).

Ridge-Trench Collision—The Southern Patagonian Cordillera Stop 4-1: Zeballos Complex and Tertiary Sediments The first stop of the day will be at the pass at ~1500 m elevation at a place called the “El Portezuelo” or sometimes referred to as Paso Zeballos. Early to middle Miocene Rio Zeballos Group fluvial sandstones and conglomerates (equivalent to the Santa Cruz Formation) that represent foreland basin

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molasse deposits are well exposed in the canyons of the Rio Zeballos and Rio Jeinimeni (Figs. 12 and 13). Conspicuous meter-wide dikes intrude the Zeballos Group in several localities. On the east side of the valley, middle Miocene to Pleistocene slab window basalts of the Meseta del Lago Buenos Aires unconformably overly the Rio Zeballos Group. The prominent

Figure 12. Geologic map of the Lago General Carrera–Buenos Aires area. Figure from Lagabrielle et al. (2004).

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Figure 13. Geologic cross section along line of section in Figure 12. Figure from Lagabrielle et al. (2004).

Monte Zeballos (2743 m), the sharp spire of Cerro Zeballos (2073 m), and an unnamed, light-colored dome (2648 m, 40 Ar/39Ar age = 3.3 ± 0.3 Ma; Brown et al., 2004) dominate the eastern skyline on the western edge of the Meseta del Lago Buenos Aires. Recent reconnaissance mapping and geochemical analyses (DaSilva et al., 2006) indicate that this a highly alkaline volcanic complex (informally named the “Zeballos Complex”) that contains rare trachyandesitic and trachydacitic lavas and intrusives that are typically of Neogene slab windowrelated plateau sequences. The Sierra Chacabuco on the west side of the valley is composed of Jurassic volcanics that are uplifted along a N-S east-vergent thrust fault that lies roughly within the axis of the valley (section B–B′ on Fig. 11). This thrust dies out to the north of Lago Buenos Aires as deformation style changes to the mild tectonic inversion characteristic of the foreland north of the Chile Triple Junction. West of Chile Chico, the main road follows the southern shore of Lago General Carrera and traverses through mostly Jurassic rhyolitic volcanics. About ~50 km west of Chile Chico, the road crosses a prominent river valley at Puerto Fachinal. On the west side of the river there are excellent roadcuts through glacial sediments and the road climbs up onto the Fachinal Moraine (Younger Dryas?). About 25 km west of Fachinal, the road cuts through the ca. 10 Ma Paso Las Llaves granite pluton. Another ~40 km west, just east of the Rio Las Duñas, outcrops of northernmost Cosmelli Basin can be seen in the low craggy hills to the south. The Cosmelli Basin (see map in Fig. 12) succession consists of late Paleocene to early middle Miocene marine and fluvial molasse sediments that are equivalent to the Centinela and Santa Cruz Formations in Argentina. The basin is deformed and shows evidence of synsedimentary contractional tectonics interpreted to be related to the same early to late Miocene defor-

mation that affected the Patagonian fold-and-thrust belt in the Argentine foreland (Flint et al., 1994). Finally, west of Puerto Guadal to Puerto Bertrand, there are numerous roadcuts of the late Paleozoic metasedimentary rocks of the Rio Lacteo Formation, which form the metamorphic basement of the southern Patagonian Cordillera. These rocks represent metamorphosed marine turbidites, limestones, and pyroclastics rocks that were intensely deformed and subjected to medium-grade greenschist facies metamorphism during pre-late Carboniferous times (Bell and Suárez, 2000). Stay overnight in cabins along the Rio Baker a few kilometers south of Puerto Bertrand. Day 5—Puerto Bertrand to Coyhaique Drive back north along the Carretera Austral (Chilean National Route 7). The late Paleozoic Rio Lacteo Formation is exposed along many roadcuts between Puerto Bertrand and Puerta Murta, a distance of ~100 km. Along this stretch of road there are many excellent viewpoints of the high peaks to the west, especially the San Valentin massif (4078 m), the highest peak in the Patagonian Cordillera. The Cretaceous Patagonian Batholith underlies these peaks, and the high topography supports the Northern Patagonian Ice Cap. There are also superb vistas looking southwest across Lago General Carrera. Near Puerto Murta there are roadcut exposures of Cretaceous granitoids of the batholith and the young (<1 Ma) Murta basalts to examine. The road continues northeastward and then west following the Rio Ibáñez to the town of Villa Castillo. The picturesque Cordón Castillo dominates the view and is made up of a roof pendant (?) of Jurassic ignimbrites overlying granitoids of the Patagonian Batholith.

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Gripp, A.E., and Gordon, R.G., 1990, Current plate velocities relative to the hotspots incorporating the NUVEL-1 global plate motion model: Geophysical Research Letters, v. 17, p. 1109–1112. Guivel, C., Lagabrielle, Y., Bourgois, J., Maury, R.C., Fourcade, S., Martin, H., and Arnaud, N., 1999, New geochemical constraints for the origin of ridge subduction-related plutonic and volcanic suites from the Chile Triple Junction (Taitao Peninsula and Site 862, LEG ODP on the Taitao Ridge): Tectonophysics, v. 311, p. 83–111, doi: 10.1016/S0040-1951(99)00160-2. Gust, D.A., Biddle, K.T., Phelps, D.W., and Uliana, M.A., 1985, Associated middle to late Jurassic volcanism and extension in southern South America: Tectonophysics, v. 116, p. 223–253, doi: 10.1016/0040-1951(85)90210-0. Halpern, M., 1973, Regional geochronology of Chile south of 50° latitude: Geological Society of America Bulletin, v. 84, p. 2407–2422, doi: 10.1130/ 0016-7606(1973)84<2407:RGOCSO>2.0.CO;2. Hervé, F., 1988, Late Paleozoic subduction and accretion in southern Chile: Episodes, v. 11, p. 183–188. Hervé, F., Aguirre, L., Godoy, E., Massone, H., Morata, D., Pankhurst, R.J., Ramirez, E., Sepulveda, V., and Willner, A., 1998, Nuevos antecedentes acerca de la edad y las condiciones P-T de los complejos metamorficos en Aysen, Chile: Buenos Aires, Actas X Congreso Latinoamericano de Geólogo, VI Congreso Nacional de Geología y Economica, v. 2, p. 134–137. Ivins, E.R., and James, T.S., 1999, Simple models for late Holocene and present-day Patagonian glacier fluctuations and predictions of a geodetically detectable isostatic response: Geophysical Journal International, v. 138-3, p. 601–624. Katz, H.R., 1972, Plate tectonics-orogenic belts in the southeast Pacific: Nature, v. 237, p. 331, doi: 10.1038/237331a0. Kay, R.W., 1978, Aleutian magnesian andesites: Melts from subducted Pacific ocean crust: Journal of Volcanology and Geothermal Research, v. 4, p. 117–132, doi: 10.1016/0377-0273(78)90032-X. Kay, R.W., and Kay, S.M., 2002, Andean adakites: Three ways to make them: Acta Petrologica Cínica, v. 18-3, p. 303–311. Kay, S.M., Ramos, V.A., Mpodozis, C., and Sruoga, P., 1989, Late Paleozoic to Jurassic silicic magmatism at the Gondwana margin: Analogy to the Middle Proterozoic in North America?: Geology, v. 17, p. 324–328, doi: 10.1130/0091-7613(1989)017<0324:LPTJSM>2.3.CO;2. Kay, S.M., Ramos, V.A., and Marquez, M., 1993, Evidence in Cerro Pampa volcanic rocks for slab-melting prior to ridge-collision in southern South America: The Journal of Geology, v. 101, p. 703–714. Kay, S.M., Ramos, V.A., and Gorring, M.L., 2002, Geochemistry of Eocene plateau basalts related to ridge collision in southern Patagonia, in Cabaleri N., Cingolani, C.A., Linares, E., López de Luchi, M.G., Ostera, H.A., and Panarello, H.O., eds.: XV Congreso Geológico Argentino Actas, v. 3, p. 60–65. Klepeis, K.A., 1994, The Magallanes and Deseado fault zones: Major segments of the South American-Scoria transform plate boundary in southernmost South America: Journal of Geophysical Research, v. 99, p. 22,001–22,014, doi: 10.1029/94JB01749. Kraemer, P.E., 1993, Perfil estructural de la Cordillera Patagonica Austral a los 50°S, Santa Cruz: XII Congreso Geológico Argentino Actas, v. 3, p. 119–125. Kraemer, P.E., 2003, Orogenic shortening and the origin of the Patagonian orocline (56° S. Lat): Journal of South American Earth Sciences, v. 15, p. 731–745, doi: 10.1016/S0895-9811(02)00132-3. Kraemer, P.E., Ploszkiewicz, J.V., and Ramos, V.A., 2002, Estructura de la Cordillera Patagonica Austral entre los 46º y 52ºS, provincial de Santa Cruz, Argentina, in Haller, M.J., ed., Geología y recursos naturales de Santa Cruz: Relatorio del XV Congreso Geológico Argentino, p. 353–364. Lagabrielle, Y., Le Moigne, J., Maury, R.C., Cotten, J., and Bourgois, J., 1994, Volcanic record of the subduction of an active spreading ridge, Taitao Peninsula (southern Chile): Geology, v. 22, p. 515–518, doi: 10.1130/ 0091-7613(1994)022<0515:VROTSO>2.3.CO;2. Lagabrielle, Y., Guivel, C., Maury, R., Bourgois, J., Fourcade, S., and Martin, H., 2000, Magmatic-tectonic effects of high thermal regime at the site of active spreading ridge subduction: The Chile Triple Junction model: Tectonophysics, v. 326, p. 255–268, doi: 10.1016/S0040-1951(00)00124-4. Lagabrielle, Y., Suárez, M., Rosello, E.A., Hérail, G., Martinod, J., Réginer, M., and de la Cruz, R., 2004, Neogene to Quaternary tectonic evolution of the Patagonian Andes at the latitude of the Chile Triple Junction: Tectonophysics, v. 385, p. 211–241. Leanza, A., 1972, Andes Patagonicos australes, in Leanza, A., ed., Geología a regional Argentina: Cordoba, Academia Nacional Ciencias, p. 689–706.

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Le Moigne, J., Lagabrielle, Y., Whitechurch, H., Girardeau, J., Bourgois, J., and Maury, R.C., 1996, Petrography and geochemistry of the ophiolitic and volcanic suites of the Taitao Peninsula-Chile triple junction area: Journal of South American Earth Sciences, v. 9, p. 43–58, doi: 10.1016/0895-9811(96)00026-0. Malumián, N., 2002, El terciario marino: Sus relaciones con el eustatismo, in Haller, M.J., ed., Geología y recursos naturales de Santa Cruz: Relatorio del XV Congreso Geológico Argentino, v. 1, p. 237- 244. Mercer, J.H., 1976, Glacial history of southernmost South America: Quaternary Research, v. 6, p. 125–166, doi: 10.1016/0033-5894(76)90047-8. Michael, P.J., 1983, Emplacement and differentiation of Miocene plutons in the foothills of the southernmost Andes [Ph.D. thesis]: New York, Columbia University, 367 p. Minster, J.B., and Jordan, T.H., 1978, Present day plate motions: Journal of Geophysical Research, v. 83, p. 5331–5354. Morata, D., Barbero, L., Suárez, M., and de la Cruz, R., 2002, Early Pliocene magmatism and high exhumation rates in the Patagonian Cordillera (46º40’S) K-Ar and fission track data: Toulouse, France, Fifth International Symposium on Andean Geodynamics (ISAG), p. 433–436. Mpodozis, C., Hervé, M., Nasi, C., Soffia, J., Forsythe, R., and Nelson, E., 1985, El magmatismo plioceno de Península Tres Montes y su relación con la evolución del Punto Triple de Chile Austral: Revista Geológica de Chile, v. 25–26, p. 13–28. Niemeyer, H., Skarmeta, J., Fuenzalida, R and Espinosa, W., 1984, Hojas Península de Taitao y Puerto Aisen, Region de Aisen del General Carlos Ibáñez del Campo: Carta Geológica de Chile, Servicio Nacional de Geología y Mineria, Chile, p. 60–61, scale: 1:500,000. Nullo, F.E., Proserpio, C., and Ramos, V.A., 1978, Estratigrafía y tectónica de la vertiente este Hielo Continental Patagónico, Argentina, Chile: VII Congreso Geológico Argentino Actas, v. 1, p. 455–470. Pankhurst, R.J., Leat, P.T., Sruoga, P., Rapela, C.W., Marquez, M., Storey, B.C., and Riley, T.R., 1998, The Chon Aike province of Patagonia and related rocks in west Antarctica: A silicic large igneous province: Journal of Volcanology and Geothermal Research, v. 81, p. 113–136, doi: 10.1016/ S0377-0273(97)00070-X. Pankhurst, R.J., Weaver, S.D., Hervé, F., and Larrondo, P., 1999, MesozoicCenozoic evolution of the North Patagonian Batholith in Aysen, southern Chile: Journal of the Geological Society, v. 156, p. 673–694. Panza, J.L., and Nullo, F.E., 1994, Mapa geológico de la provincia de Santa Cruz, Republica de Argentina: Buenos Aires, Servicio Geológico de Argentina, v. 1, scale: 1:750,000. Pardo-Casas, F., and Molnar, P., 1987, Relative motion of the Nazca (Farallon) and South American Plates since Late Cretaceous time: Tectonics, v. 6, p. 233–248. Petford, N., and Turner, P., 1996, Reconnaissance 40Ar/39Ar age and paleomagnetic study of igneous rocks around Coyhaique, S. Chile: Saint Malo, France, Third International Symposium on Andean Geology (ISAG), p. 625–627. Pittion, J.-L., and Gouadain, J., 1992, Source rocks and oil generation in the Austral Basin: Buenos Aires, Proceedings of Thirteenth World Petroleum Congress, p. 113–120. Ramos, V.A., 1988, Late Proterozoic-Early Paleozoic of South America: A collisional history: Episodes, v. 11, p. 168–174. Ramos, V.A., 1989, Foothills structure in Northern Magallanes Basin, Argentina: AAPG Bulletin, v. 73, p. 887–903. Ramos, V.A., 2005, Seismic ridge subduction and topography: Foreland deformation in the Patagonian Andes: Tectonophysics, v. 399, p. 73–86, doi: 10.1016/j.tecto.2004.12.016. Ramos, V.A., and Kay, S.M., 1992, Southern Patagonian plateau basalts and deformation: Backarc testimony of ridge collision: Tectonophysics, v. 205, p. 261–282, doi: 10.1016/0040-1951(92)90430-E.

Ramos, V.A., Kay, S.M., and Singer, B.S., 2004, Las adakitas de la Cordillera Patagónica: Nuevas evidencias geoquímicas y geocronológicas: Revista de la Associación Geológica Argentina, v. 59, p. 693–706. Riccardi, A.C., 1971, Estratigrafía en el oriente de la Bahía de la Lancha, Lago San Martin, Santa Cruz, Argentina: Museo de la Plata Revista (Geología), v. 7, p. 245–318. Riccardi, A.C., 1988, The Cretaceous system of southern South America: Geological Society of America Memoirs 168, 161 p. Singer, B.S., Ackert, R.P., Jr., and Guillou, H., 2004, 40Ar/39Ar and K-Ar chronology of Pleistocene glaciations in Patagonia: Geological Society of America Bulletin, v. 116, p. 434–450, doi: 10.1130/B25177.1. Skarmeta, J., and Castelli, J.C., 1997, Intrusión sintectónica del Granito de Las Torres del Paine, Andes Patagónicos de Chile: Revista Geológica de Chile, v. 24, p. 55–74. Stern, C.R., and Kilian, R., 1996, Role of the subducted slab, mantle wedge, and continental crust in the generation of adakites from the Andean Austral Volcanic Zone: Contributions to Mineralogy and Petrology, v. 123, p. 263–281, doi: 10.1007/s004100050155. Stern, C.R., Frey, F.A., Futa, K., Zartman, R.E., Peng, Z., and Kyser, T.K., 1990, Trace element and Sr, Nd, Pb, and O isotopic composition of Pliocene and Quaternary alkali basalts of the Patagonian Plateau lavas of southernmost South America: Contributions to Mineralogy and Petrology, v. 104, p. 294–308, doi: 10.1007/BF00321486. Suárez, M., and de la Cruz, R., 2001, Jurassic to Miocene K-Ar dates from eastern central Patagonian Cordillera plutons, Chile (45°–48°S): Geological Magazine, v. 138, p. 53–66, doi: 10.1017/S0016756801004903. Suárez, M., and Pettigrew, T.H., 1976, An upper Mesozoic island arc-back-arc system in the southern Andes and South Georgia: Geological Magazine, v. 113, p. 305–328. Suárez, M., de la Cruz, R., and Bell, C.M., 2000, Timing and origin of deformation along the Patagonian fold and thrust belt: Geological Magazine, v. 137, p. 345–353, doi: 10.1017/S0016756800004192. Thomas, C.R., 1949, Geology and petroleum exploration in the Magallanes Province, Chile: AAPG Bulletin, v. 33, p. 1553–1578. Thomson, S.N., Hervé, F., and Stockhert, B., 2001, Mesozoic-Cenozoic denudation history of the Patagonian Andes (southern Chile) and its correlation to different subduction processes: Tectonics, v. 20, p. 693–711, doi: 10.1029/2001TC900013. Thorkelson, D.J., 1996, Subduction of diverging plates and the principles of slab window formation: Tectonophysics, v. 255, p. 47–63, doi: 10.1016/0040-1951(95)00106-9. Welkner, D., 1999, Geología del area del Cerro de San Lorenzo: Cordillera Patagónica oriental, XI Región de Aysén, Chile (47°25′–47°50′S) [M.S. thesis]: Santiago, Departamento de Geología, Universidad de Chile, 148 p. Welkner, D.R., 2000, Geocronología de los plutones del área de del Cerro San Lorenzo, XI Región de Aysén: IX Congreso Geológica de Chile (Puerto Varas) Actas, v. 2, p. 269–273. Winslow, M.A., 1981, Mechanisms for basement shortening in the Andean foreland fold belt of southern South America, in McClay, K-.R., and Price, N.J., eds., Thrust and nappe tectonics: Geological Society of London, Blackwell Scientific Publications, p. 513–528. Winslow, M.A., 1982, The structural evolution of the Magallanes Basin and neotectonics in the southernmost Andes, in Cradock, C., ed., Antarctic geoscience: Madison, University of Wisconsin, p. 143–154. Yogodzinski, G.M., Lees, J.M., Churikova, T.G., Dorendorf, F., Woerner, G., and Volynets, O.N., 2001, Geochemical evidence for the melting of subducting oceanic lithosphere at plate edges: Nature, v. 409, p. 500–504, doi: 10.1038/35054039. MANUSCRIPT ACCEPTED BY THE SOCIETY 10 JANUARY 2008

Printed in the USA

The Geological Society of America Field Guide 13 2008

Field trip guide: Andean Cordillera and backarc of the south-central Andes (~38.5°S to 37°S) Tomás Zapata Gonzalo Zamora Valcarce Repsol-YPF, Talero 360, Neuquén 8300, Argentina Andrés Folguera Laboratorio de Tectónica Andina, Universidad de Buenos Aires, Consejo Nacional de Investigaciones Científicas y Técnicas (CONICET), Buenos Aires 1428, Argentina Daniel Yagupsky Laboratorio de Tectónica Andina, Universidad de Buenos Aires, Buenos Aires 1428, Argentina

ABSTRACT The Andes of the Neuquén Mesozoic basin have experienced multiple episodic tectonic events as a consequence of the changes of the plate tectonic boundary configuration. Each episode of deformation has overprinted the previous one, making it difficult to unravel the Andean tectonic history. The first deformation event took place in the uppermost Cretaceous with the formation of the Agrio fold-and-thrust belt. This event was related to the shallowing of the subducting plate recorded by the migration of the volcanic arc toward the foreland. During the late Oligocene–early Miocene, an extensional event, related to the steepening of the subducted plate, affected only the hinterland region causing the opening of the Cura Mallín basin. This basin was closed during the late Miocene, together with the development of a new fold-and-thrust belt that reactivated the previous structures. During the Late Tertiary, two more episodes of extension and compression affected the Andean area. The trip focuses on the field evidence that documents this complex history of evolution by looking at evidence of the sequence of the distinct tectonic events. Keywords: Andes, Neuquén, tectonic events, shallowing.

*E-mails: [email protected]; [email protected]; [email protected]; [email protected]

Zapata, T., Zamora Valcarce, G., Folguera, A., and Yagupsky, D., 2008, Field trip guide: Andean Cordillera and backarc of the south-central Andes (~38.5°S to 37°S), in Kay, S.M., and Ramos, V.A., eds., Field trip guides to the Backbone of the Americas in the southern and central Andes: Ridge collision, shallow subduction, and plateau uplift: Geological Society of America Field Guide 13, p. 23–55, doi: 10.1130/2008.0013(02). For permission to copy, contact editing@ geosociety.org. ©2008 The Geological Society of America. All rights reserved.

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INTRODUCTION The field trip covers the westernmost part of the Neuquén province in Argentina (Fig. 1). Geologically, the trip takes place in the fold-and-thrust belt of the Neuquén Basin between 36° and 38°S on the eastern slope of the Andes. The aim of the trip is to visit key localities where observations have been made in recent years that have improved our understanding of the history of the region. The area has undergone a complex tectonic evolution that is suggested to be linked to changes in the dip of the underlying subduction zone. These slab dip changes have led to the migration and retreat of the volcanic arc and the deformation front. The field guide focuses on the evidence for the Mesozoic to Recent deformational events that affected this part of the Andes.

volcanic pulse that is broadly exhibited in the northern Patagonian Andes. Nonmarine sedimentary facies interbedded with these volcanic rocks were correlated with Late Jurassic to Early Cretaceous marine sequences of the Neuquén Basin (Zanettini et al., 1987). Gutiérrez and Miniti (1985) could be credited with the first high-quality study of the Cenozoic Cura Mallín basin within the study area. At the same time, a considerable number of radiometric ages were obtained on the western slope of the Andes between 36° and 38°S in Chile. These ages show that the volcanic units and associated sedimentary facies in Chile accumulated in a late Oligocene–early Miocene basin (Niemeyer and Muñoz, 1983). The studies of Burns and Jordan (1999) and Jordan et al. (2001) subsequently showed that the sequences on the eastern slope of the Andes were of similar Tertiary age, and have been tectonically inverted more recently.

Brief Review of the History of Geologic Exploration and Research

GEOLOGICAL SETTING

The first geologic descriptions of the region by Bodenbender in 1891 were of the volcanic rocks of the Cerro Tromen region, which is located just east of the city of Chos Malal (Fig. 1). Evidence for a young eruption was reported. The first geologic transect through this part of the Andes was made by Burckhardt in 1900 (Fig. 2), who crossed the Andes through the Pino Hachado Pass and described the main morphostructural units. In the first half of the last century, a large number of geologists studied the sedimentary sequences (e.g., Gerth, 1928; Weaver, 1931; Groeber, 1929, 1946a, 1946b). In 1946, Herrero Ducloux produced a synthesis of the geology of Neuquén. During the 1970s and 1980s, the area was intensively explored by YPF geologists (the former Argentine state petroleum company) leading to advancements in the understanding of the stratigraphic and structural framework of the region (e.g., Gulisano et al., 1984; Legarreta and Gulisano, 1989). Ramos (1978, 1981, 1998) described the structure of the Agrio foldand-thrust belt and showed that the belt consisted of an inner western sector characterized by basement inversion and an eastern outer sector characterized by thin-skinned deformation. Llambías and Malvicini (1978) and Llambías and Rapela (1987, 1989) studied the igneous rocks. They dated and described the Collipilli igneous rocks, associating them with different volcanic provinces. Recent studies (Zapata et al., 2002; Zapata and Folguera, 2005; Zamora Valcarce et al., 2006) demonstrate structural inversion of the basement at the frontal part of the fold-and-thrust belt. Recent 40Ar/39Ar ages have shown that this deformation has a minimum age of Cretaceous. The transect of the main cordillera covered by this field guide has been studied more recently. In the 1960s to the 1980s, there was a tendency for researchers working in the area to correlate the sedimentary and volcaniclastic sequences in the Andes of Neuquén with units outcropping to the east in the Agrio fold-andthrust belt and to the south along the northern Patagonian Andes. Researchers such as Pesce (1981) mistakenly assigned Paleogene volcaniclastic units in the Neuquén Andes to a Middle Cretaceous

The Andes in this region (36°–38°S latitude) show distinctive deformational characteristics that result from alternation of periods of generalized extension followed by periods of compression. The Agrio fold-and-thrust belt in the eastern part reflects these processes (Fig. 3). Deformation in the Agrio belt started during the Cretaceous and was reactivated in the late middle to late Miocene. The Loncopué trough to the west of the Agrio belt (Fig. 3) is a long depression in the foothills that parallel the Principal Cordillera. The Loncopué trough consists of a complex half-graben system that was produced during the Oligocene and extensionally reactivated in the Pliocene-Pleistocene. The northern part of the Loncopué trough is currently located in the contractional orogenic front. The nature and volume of arc-related igneous rocks, the location of the volcanic fronts, expansions and retreats of the magmatism, and the associated igneous activity in the foreland and superimposed structural styles provide evidence for the alternating tectonic regimes. The various Jurassic to Recent stages are correlated with changes in the geometry of the Benioff zone. Periods of subduction-zone steepening are associated with large volumes of poorly evolved magmas and generalized extension, whereas shallowing of the subduction zone is linked to foreland migration of more evolved magmas associated with contraction and uplift in the Principal Cordillera. The purpose of this field guide is to show, in chronological order, the evidence for the complex evolutionary history of the Neuquén Andes at this latitude. The first day of the trip travels from Neuquén city to Chos Malal with a few stops to recognize the structural styles. The second and third days are in the Agrio fold-and-thrust belt looking at the evidence for the Cretaceous and Miocene deformation. The fourth day is in the region west of the Agrio fold-and-thrust belt in the foothills of the modern orogenic front. The stops on this day focus on the late Oligocene to early Miocene extensional collapse of the inner sectors of the Agrio fold-and-thrust belt, and the subsequent late Miocene compressional inversion.

60°W

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BOLIVIA

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BRASIL

San Rafael

URUGUAY

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PA RA GU AY

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. Malargüe !

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Ahuca Mahuida Volcano

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Kilometers 71°W

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Figure 1. Location map of the region covered in the field guide.

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Figure 2. First transect through the cordillera at 39°S from Burckhardt (1900) shows the names of the main morphostructural units, which are still in use.

Neuquén Basin The Neuquén Basin is a triangular-shaped basin that contains a Mesozoic-Cenozoic sedimentary succession that is at least 7 km thick. The tectonic evolution of the basin records a continuous subsidence lasting at least 220 million years that extends from the Early Jurassic to the end of the Early Cretaceous. During this time the region was a marine retroarc basin (Ramos, 1989; Vergani et al., 1995). A stratigraphic column summarizing the major units in the region is shown in Figure 4. As a result of the Andean deformation that progressed from west to east, the basin became a foreland basin at the beginning of the Late Cretaceous. Vergani et al. (1995) summarized the main tectonic elements of the basin (Fig. 5) that include a north-south fold-and-thrust belt in the west (in the region of the field trip), the east-west–running Huincul Arch, and an inner basin separated from the fold-andthrust belt by a structural high that corresponds to the Chihuidos High and a frontal syncline. The basement (pre-Jurassic) underlying the Neuquén Basin is composed of a rhyolite-ignimbrite suite (Choiyoi Group) of enormous proportions associated with extensional tectonics and molasse deposition (Ramos and Kay, 1992). Uliana and Legarreta (1993) associated these sequences with the extensional collapse of a Permian-Triassic orogenic belt. From the Late Triassic through the Early Jurassic, basement-involved extensional faults created a series of half grabens (Vergani et al., 1995) that formed the initial Neuquén

Basin. During this stage, the pre-Cuyo and Cuyo Groups were deposited in the developing half grabens (Fig. 4). This period of rifting ended with the deposition of the Tábanos Formation evaporites. During the Late Jurassic, general uplift and erosion affected much of the Neuquén Basin and marked a conspicuous change in the sedimentation style from clastic and volcaniclastic deposits to the evaporite-limestone–dominated deposits of the Lotena Group. The end of this phase coincides with a period of regional subsidence that Vergani et al. (1995) attributed to compressional relaxation. Later deposition of the Mendoza Group (Fig. 4) in the Andico cycle reflects marine encroachment and expansion of the basin (Gulisano et al., 1984). The Mendoza Group starts with fluvial sandstones of the Tordillo Formation, which are overlain by marine shales of the Vaca Muerta Formation that are the most prolific source rock of the basin. This cycle also shows structural inversion events that interrupted marine sedimentation and caused the erosion of older units in the southern and eastern margins of the basin (Vergani et al., 1995). Next is the Mulichinco Formation. Tectonic quiescence resulted in renewed marine transgression and deposition of the Agrio Formation. General shallowing of the basin and evaporitic and clastic sedimentation (Rayoso Group, Fig. 4) marks the culmination of the Andico cycle. Renewed tectonic activity and inversion during the Early Cenomanian reactivated the provenance regions and resulted in accumulation of the continental deposits of the Neuquén Group. This sedimentation persisted until the end of

Andean Cordillera and backarc of the south-central Andes 72°W

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Quaternary retro-arc basalts Liquiñe-Ofqui Fault Zone

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Figure 3. Main geological provinces of the Neuquén Andes at this latitude. Vn.—volcano.

the Cretaceous, when the continental and marine deposits of the Malargüe Group marked the culmination of the Riográndico Cycle (Fig. 4) (Legarreta and Gulisano, 1989). In late Cretaceous, the Neuquén Basin entered a new phase of deformation and basin subsidence. The north-trending Andean fold-and-thrust belt encroached on the western margin of the basin deforming the Mesozoic prism in front of it. Preexisting structures, as well as the mechanical properties of the Mesozoic sequence, controlled the deformation style in this belt.

Agrio Fold-and-Thrust Belt The Agrio fold-and-thrust belt is located south of the Cortaderas Lineament. The belt is bounded to the east by the Los Chihuidos high and to the west by the Loncopué trough. The structure of the Agrio fold-and-thrust belt is characterized by a combination of thin-skinned and thick-skinned structures (Fig. 6), which show good examples of detachment folds. The geologic sections show north-northwest–trending

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Main decollments in the fold and thrust belt Synrift sequences inverted into the fold and thrust belt Not inverted synrift sequences

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Figure 4. Stratigraphic column showing the major sedimentary and volcanic units of the area. After Zapata and Folguera (2005).

Andean Cordillera and backarc of the south-central Andes 69°W

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External Fold and Thrust Belt Huincul Arch

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Figure 17 Andacollo . ! Chos Malal

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Rincon de los Sauces

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Figure 6

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Figure 5. Tectonic elements of the Neuquén Basin after Vergani et al. (1995).

structures with a difference in structural style between the western and eastern parts. To the west, the structure is characterized by thick-skinned structures with broad anticlines, which are the product of basement inversion (Ramos, 1998). To the east, the Agrio fold-and-thrust belt is composed of large doubly-plunging anticlines detached in the Auquilco evaporites. The anticlines are separated by broad synclinoria with general rhombic shapes that

reflect the existence of basement blocks at depth (Ramos, 1978; Viñes, 1985; Zapata et al., 1999; Zapata et al., 2002). Due to the differences in structural style, the Agrio fold-and-thrust belt has been divided into two regions (Fig. 6). The western part or inner sector exposes folds related to the inversion of a Mesozoic extensional structure called the Tres Chorros extensional system (Vergani et al., 1995). The eastern or outer sector (Ramos, 1977;

Auquilco Formation Choiyoi Group

Upper Mendoza Group

Lower Mendoza Group Cuyo Group

Rayoso Group

Jurassic Detachment –15

–10

Figure 6. Cross section through the inner and the outer zones (sectors) of the Agrio fold-and-thrust belt showing the differences in structural style (modified from Zapata and Folguera, 2005; Zamora Valcarce et al., 2006).

Synorogenic deposits

Neuquén Group

Figure 21 PMu.x-10

CLMu-X.1 PiMu-X.1 PDS-X.1

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Figure 7 CMo-X.1

–5

The stops on Day 1 provide an introduction to the stratigraphy and structural style of the external zone of the Agrio foldand-thrust belt in this region of the Andes (Fig. 8).

0

DAY 1—TRIP FROM THE CITY OF NEUQUÉN TO CHOS MALAL

Inner zone

Structure of the Outer Sector The structure of the Outer Sector is composed of thinskinned tight folds associated with deep faults. Borehole and twodimensional seismic data show that the deep faults are detached from the Jurassic Auquilco evaporites and propagate up through the Mesozoic sequence (lower and upper Mendoza Group) until they reach the Cretaceous Huitrín evaporites (Rayoso Group) where they form fault-bend fold structures (Fig. 6). The upper stratigraphic units of these structures, which correspond to the Agrio Formation (upper Mendoza Group), have been locally deformed by flexural folding adding a detachment folding component (Zapata et al., 2002; Zamora Valcarce, 2007). Seismic information also shows that the external structure of the Agrio fold-andthrust belt is characterized by a refolded triangle zone bounded on the eastern side by a backthrust that produces a fault-related fold geometry (Zapata et al., 2002; Zamora Valcarce, 2007).

Chorriaca Fold Type

Structure of the Inner Sector The Inner Sector of the Agrio fold-and-thrust belt includes the southern extension of the Cordillera del Viento basement uplift. This uplift belongs to the Tres Chorros extensional system, which is composed of northwest-trending, inverted half grabens forming broad anticlines. One structure is the Cerro Mocho anticline, a doubly vergent basement uplift that resembles a “pop-up” structure. Two-dimensional seismic data (Fig. 7) show that the eastern limb of the structure is affected by a deep fault that cuts across the sedimentary sequence through the Auquilco Jurassic evaporites and transfers >6 km of shortening to the thin-skinned structures of the Outer Sector (Zapata et al., 2002). An anomalous thickness of Jurassic synrift sequences recorded in borehole data is interpreted to be associated with an extensional half graben that was inverted during the Andean orogeny.

Chihuidos High

Ramos and Barbieri, 1989) is composed of tight, axially extended anticlines that bound deep basement blocks (Zapata et al., 2002). This part of the fold-and-thrust belt has experienced several episodes of deformation from the late Cretaceous to the Miocene (Zapata et al., 2002, 2005; Cobbold and Rosello, 2003; Zamora Valcarce et al., 2006), which are recorded by volcanic rocks and synorogenic deposits. In the last compressional pulse, which could have begun in the middle Miocene, the whole fold-andthrust belt was thrust toward the foreland. The thrusts probably used the preexisting Jurassic detachment (Zapata et al., 2002) that was inherited from the extensional period of the Neuquén Basin. The related Miocene synorogenic deposits of the Agrio fold-and-thrust belt are buried in the Bajo de Añelo area to the east of the field trip region (Ramos, 1999).

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Choiyoi Group Figure 7. Two-dimensional seismic line through the Cerro Mocho anticline showing the insertion of the basement (Choiyoi Group) into the sedimentary cover and the transfer of shortening into the foreland. TWT—two-way time in seconds. After Zamora Valcarce et al. (2006).

Stop 1: Rio Agrio Anticline Location: Agrio River near town of Bajada del Agrio at 38°20′47″S; 69°58′3″W; 607 m. The Agrio anticline is located at the thrust front zone of the southern termination of the Agrio fold-and-thrust belt. At

first glance, the geometry can be described as a simple, doubly plunging structure (Fig. 9) with a north-south axis characterized by a smooth crest and steeply dipping flanks with dips up to 40°–60°. The anticlinal limbs are part of broad synclines that separate the Agrio anticline from the unexposed Quili Malal– Esquinero anticline to the east and Cordón del Salado anticline

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Figure 8. Image showing the geographic features and stops for Day 1 of the trip.

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Figure 9. Photo looking to the north showing a panoramic view of the Rio Agrio anticline.

to the west. The core of the anticline, where the Agrio Formation crops out, is affected by a series of strike-slip and minor thrust faults that offset the key bed horizon of the Avilé Member. These structures are interpreted as part of the crestal extension due to the buckling process of folding. This is also an historical stop, since the first well of the Agrio fold-and-thrust belt, the RA.x-1, was drilled here in 1935 by the Standard Oil Company. The geological studies that led to the drilling proposal were based on the presence of oil seeps, following the classical “anticline” play (surface anticline with associated oil seeps). Later, in 1949, YPF (the former Argentine national oil company) re-drilled the play reaching a final depth of 3220 m below ground level and ending in the Jurassic Tordillo Formation. The results in both cases were dry holes.

décollement horizons producing internal disharmonies (Fig. 10). The anticline has steeply dipping to overturned flanks, which are affected by minor thrusting and back thrusts that accommodate the internal layer-parallel shear deformation. Stop 3: Pampa del Salado, Road to Chos Malal Location: 38°01′53″S; 70°02′54″W; 993 m. The purpose of Stop 3 is to review the Lower Cretaceous stratigraphy of the upper section of the Agrio Formation and the individual members of the Huitrín Formation. The Huitrín Formation is divided into three members: (1) Lower Troncoso, which is composed of a lower fluvial and an upper eolian section; (2) Upper Troncoso, which is constituted by evaporites; and (3) La Tosca Member, which is formed by interbedded limestones and shales.

Stop 2: Cordón del Salado Location: 38°16′10″S; 70°03′28″W; 758 m. The Cerro La Mula–Naunanuco anticline, which is the longest structure of the Agrio fold-and-thrust belt, extends for more than 60 km. The structure is an axially extended, doubly plunging anticline, whose axis is broken by lateral ramps that may accommodate local differences in shortening, acting as transfer zones. The oldest unit exposed at the core of the anticline is the Vaca Muerta Formation. The outcrop pattern shows that these are geometrically complex structures that result from multiple

DAY 2—CRETACEOUS DEFORMATION AND THE VOLCANIC ARC For the next three days, evidence for the different deformational events affecting the Neuquén Andes at this latitude will be examined (Fig. 11). The stops on Day 2 focus on the inner part of the Agrio fold-and-thrust belt where evidence for the Cretaceous deformation event is displayed. Evidence for the timing of this deformation comes from 40Ar/39Ar ages (Zamora Valcarce et al., 2006), crosscutting relationships, and paleomagnetic data.

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Figure 10. Photo and cross section looking to the north showing the southern plunge of the La Mula– Naunauco anticline. The section illustrates the complex structural style of the Agrio fold-and-thrust belt. Interactions between anticlines have developed triangle zones at the front of the outer sector. After Zamora Valcarce (2007).

This region is partially covered by the Upper Cretaceous–Paleocene Collipilli volcanic field. Magmatic rocks in the Collipilli consist of volcanic facies that are associated with cones, lava flows, dikes, sills, and laccoliths. Llambías and Malvicini (1978) were the first to describe these volcanic rocks. Later, Llambías and Rapela (1987) included them in the Neuquén-Mendoza Volcanic Province, which encompasses the units between 38°30′S and 34°S (Groeber, 1946a, 1946b; Yrigoyen, 1972; Bettini, 1982; Kozlowski et al., 1987; Haller et al., 1985). The intrusive series of the Collipilli Formation was emplaced as laccoliths and associated sills. These laccoliths intruded the contact between the Agrio and the Rayoso Formations, filling the spaces left by the evaporites of the Huitrín Formation during the folding process (Llambías and Malvicini, 1978; Llambías and Rapela, 1989). The volcanic rocks in the Collipilli area include extrusive domes, different types of breccias and volcanic agglomerates, pyroclastic flow deposits, and massive lava flows. These volcanic rocks unconformably cover the Agrio, Huitrín, and Rayoso Formations that are folded in the Collipilli syncline. Therefore, the area was already uplifted (and partially eroded) before or at least simultaneously with the magmatic event (Zamora Valcarce et al., 2006). Llambías and Rapela (1987, 1989) used geochemical analyses and K/Ar whole-rock ages to correlate the Collipilli region magmatic rocks with Paleogene units mapped in the Molle Formation in the Andean Cordillera. They proposed subdividing the Molle Group into two formations: (1) a subvolcanic facies called the Collipilli Formation, with ages ranging from 50 to 45 Ma (early and middle Eocene) (Llambías and Rapela, 1989), and (2) a middle Eocene volcanic facies called the Cayanta Formation (Rapela and Llambías, 1985) with one K/Ar age of 39 ± 9 Ma (Llambías and Rapela, 1989). They noted that the volcanic rocks mapped in the Molle Formation

to the north had yielded K/Ar whole-rock ages of 71.5 ± 5 Ma (Llambías et al., 1978), showing that not all of the volcanic rocks in the Molle Formation could be of Eocene age. Recently, Zamora Valcarce et al. (2006) used new 40Ar/39Ar ages ranging from 73 to 65 Ma from the Collipilli volcanic sections to reassign these magmatic rocks to the end of the Cretaceous. These authors also found evidence for a previous magmatic event, not previously recorded, with 40Ar/39Ar ages of 100 Ma (see below). Geochemical studies (Zamora Valcarce et al., 2006) show a strong arc to backarc signature (La/Ta = 66; Ba/La = 21; Ta/Hf = 0.10) and relatively flat rare earth element (REE) patterns (La/Yb = 10; La/Sm = 5.5; Sm/Yb = 2.2). The subvolcanic rocks of the Cerro Naunauco area are less evolved than those from the Collipilli region. Even though the igneous rocks from Collipilli and the Cerro Naunauco localities show different degrees of differentiation, the trace elements and incompatible rare earths indicate a similar source. Stop 1: Cerro Naunauco Laccolith Location: Near the town of Naunauco at 37°38′44″S; 70°10′11″W; 1100 m. The Cerro Naunauco is a laccolith that intruded into the northern termination of the La Mula–Naunauco anticline (Fig. 11) between the contact of the Agrio Formation and Huitrín Formation (Fig. 12). The rocks at this stop show an unconformable relationship between the igneous units of the Naunauco laccolith and the sedimentary rocks of the Mendoza Group. The new 40Ar/39Ar dates recorded for the Naunauco laccolith (Zamora Valcarce et al., 2006; Zamora Valcarce, 2007) have given an age of 65.50 ± 0.46 Ma. The crosscutting relationship puts an Upper Cretaceous age constraint on this deformational event.

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Figure 11. Image showing the geographic features and stops for Days 2 and 3 of the trip. This part of the trip is focused in the inner part of the Agrio fold-and-thrust belt, where there is evidence of both Cretaceous and Miocene deformation.

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Figure 12. Photographs of igneous rocks of the Cerro Naunauco laccolith unconformably overlying the deformed sedimentary sequence.

Stop 2: Sills of the Collipilli Formation

Stop 3: Synorogenic Tralalhué Conglomerate

Location: East of the town of Tralalhué at 37°39′14″S; 70°14′47″W; 1068 m. A sill dated at 56 Ma cuts the backlimb of the La Mula– Naunauco anticline at this stop (Zamora Valcarce et al., 2006). This sill, which belongs to the Collipilli Formation, intrudes the Agrio Formation that dips ~65° to the west on this limb of the structure. Paleomagnetic studies permit the structural evolution at the time of the sill intrusion to be reconstructed (Fig. 13). When the sill is restored to horizontal, the attitude of the paleopole differs from the regional Paleocene pole (Fig. 13). To match this difference, the sedimentary sequence needs to be tilted 25° to the west to restore its position prior to the intrusion (Fig. 13). This shows that the backlimb of the La Mula–Naunauco structure already existed in the Paleocene at 56 Ma, and that the sill intruded an already deformed sequence (Zamora Valcarce, 2007; Zamora Valcarce et al., 2007). A subsequent event tilted the stratigraphic sequence, together with the sill, to the present 65° dip.

Location: East of the town of Tralalhué at 37°37′53″S; 70°16′45″W; 1113 m. The Tralalhué conglomerate was first described by Ramos (1998) to describe conglomerate deposits on the western flank of the Cerro Naunauco anticline. These deposits have andesite, sandstone, and limestone clasts. Internally, the sequence is composed of several 1–2 m thick cycles that both thicken and thin upward, and show an upward decrease in grain size. The cycles have erosive bases with stacked channel patterns. The conglomerates overlie an angular unconformity of 18°–20° on the Naunauco igneous rocks and onlap the Rayoso Formation (Fig. 14). Restoring the Tralalhué conglomerate to horizontal shows that the Rayoso Formation was deformed at the time of Tralalhué deposition (Zamora Valcarce, 2007; Zamora Valcarce et al., 2007). These observations are in accord with those at Stop 2 (Fig. 13) where the Collipilli sills intrude a deformed sedimentary sequence (Agrio Formation) with a

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Figure 13. (A) Paleomagnetic pole for site P1-6 in situ with 100% and 66% structural corrections compared with the Upper Cretaceous pole. (B) Photo of one of the sills to the south. Field data show that the sills were intruded into a deformed sedimentary sequence and later tilted to their present position. After Zamora Valcarce et al. (2007).

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Figure 14. Schematic sketch of the Tralalhué piggy-back basin and photographs showing the unconformity between the Tralalhué conglomerate and the Rayoso, Agrio, and Collipilli Formations (Zamora Valcarce, 2007).

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similar dip. Farther west, the Tralalhué deposits dip 8°–12° west, and in the western border of the piggy-back basin, the dip is up to 70° east. A late Miocene age for the Tralalhué conglomerate is based on mammal fossil remains (Repol et al., 2002). The clastic composition of the conglomerate records an unroofing sequence related to the uplift of the La Mula–Naunauco anticline.

Stop 5: Extrusive Rocks Location: Near the town of Collipilli at 37°45′37″S; 70°19′26″W; 1108 m. A typical volcanic succession of the Collipilli Formation (as defined at Stop 2) is exposed at Stop 5 (Fig. 15). Walking from the road to the small cone in a southeast direction, a series of fluid lava flows with intraclasts can be observed. A few volcanic bombs and clastic lava flows are exposed in the middle part of the section. Proximal volcanic agglomerate crops out around the extrusive dome.

Stop 4: Panoramic View of the Collipilli Volcanic Arc Location: Near the town of Collipilli at 37°45′37″S; 70°19′26″W; 1108 m. Stop 4 shows a panoramic view of the expanse of the Collipilli volcanic field. The main components of the Collipilli volcanic field are extrusive domes and laccoliths like the Cerro del León. 40Ar/39Ar ages (Zamora Valcarce et al., 2006) indicate these rocks are upper Cretaceous to lower Paleocene. The Collipilli volcanic field unconformably overlies Upper Cretaceous units that form a broad synclinorium called the Collipilli syncline. The ages and crosscutting relationships impose a minimum age of Late Cretaceous for one of the deformational events. The geochemical signatures of the volcanic rocks show clear arc-like characteristics (Zamora Valcarce et al., 2006) (La/Ta = 49; Ba/La = 223; Ta/Hf = 0.10) and relatively flat REE patterns (La/Yb = 12.2; La/Sm = 2.26; Sm/Yb = 5.39).

Stop 6: Unconformity of the Collipilli Formation (as Defined at Stop 2) Location: Next to the town of Collipilli at 37°46′19″S; 70°20′49″W; 1145 m. This stop on the western limb of the Collipilli synclinorium shows the unconformity between the Collipilli Formation (as defined at Stop 2) and the Mendoza Group. To the north, the magmatic rocks overlie the Huitrín Formation, whereas to the south, they overlie the Huitrín and the Rayoso Formations (Fig. 15). The Collipilli igneous rocks are restricted to the syncline. The difference in erosional characteristics has produced an inversion of relief in the region because the syncline is at a higher elevation than the anticline.

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Figure 15. Photograph and sketches showing the unconformity between the Collipilli Group volcanic rocks and the Bajada del Agrio Group (Rayoso, Huitrin, and Agrio Formations; see Fig. 4). To the north, the igneous rocks overlie the Huitrín Formation and to the south the Rayoso Formation.

Andean Cordillera and backarc of the south-central Andes Stop 7: Cerro Mocho Dikes Location: Two stops close to each other. The first is next to Cerro Mocho, east of Coihuico at 38°04′39″S; 70°14′01″W; 1241 m. The other is at 38°09′33″S; 70°09′08″W; 849 m. Leanza and Hugo (2001) mapped the dikes in the Cerro Mocho region as a subunit of the Collipilli Formation. This series of east-west–trending dikes was emplaced along preexisting structures. The most distinctive is a 19-km–long, east-west dike that cuts the Cerro Mocho anticline (Fig. 16). This dike is better seen on the Landsat image than in the field due its lack of topographic expression. In detail, the dike is not a single unit, but rather a series of dike segments with chilled margins. Repol et al. (2002) mapped the Cerro Mocho dikes as the Pichaihue andesite, and assigned them a Miocene age based on field relationships and correlations with other Miocene volcanic rocks in adjacent regions (Rovere and Rossello, 2001). However, two single-crystal 40Ar/39Ar analyses have yielded ages of 101.99 ± 0.69 Ma and 90.00 ± 4.06 Ma (late Early Cretaceous– Albian; Zamora Valcarce et al., 2006). These ages are older than those obtained in the Collipilli and Naunauco igneous rocks and can be interpreted as being related to a 90–100 Ma magmatic event, not previously reported in this region.

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As discussed by Zamora Valcarce et al. (2006), the Cerro Mocho dikes are subalkaline basalts with relatively high Nb/Y ratios that vary between 0.3 and 0.4 indicating only a weak convergent continental margin signature. Their small Nb anomalies with respect to Th and Ce and high Ti and Y contents are transitional between those of arc and mid-oceanic ridge basalts. Both the high heavy rare earth and incompatible element contents require a distinct origin for the Cerro Mocho dikes compared to the arc-like Naunauco and Collipilli magmas. As such, both the geochemistry and age of the Cerro Mocho volcanic rocks are distinct from those of the Collipilli Formation (as defined at Stop 2) (Zamora Valcarce et al., 2006). DAY 3—MIOCENE DEFORMATION AT THE FRONTAL PART OF THE AGRIO FOLD-AND-THRUST BELT The stops on Day 2 were focused on the inner sector of the Agrio fold-and-thrust belt where it was possible to see clear evidence for the Cretaceous deformation that affected this part of the Andes. The only evidence for Miocene reactivation was the Tralalhué synorogenic deposits and paleomagnetic data. The stops on Day 3 focus on the outer sector of the Agrio fold-and-thrust belt (Fig. 11) where the structural deformational

Figure 16. Photographs of different views of the Lower to Middle Cretaceous Cerro Mocho dikes that are intruded into sedimentary strata as young as the Avilé Member of the Agrio Formation (see Fig. 4).

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Figure 17. Regional cross section through the Cordillera del Viento and the Tromen massif, reaching the eastern limit of the fold-and-thrust belt (Zapata et al., 1999).

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Locations: View at two stops close to each other—one east of the town of Chos Malal, south of Tromen Volcano (Route 52) at 37°20′03″S; 70°08′35″W; 1342 m; the second near Laguna Auquilco (Route 52) at 37°21′29″S; 69°59′39″W; 1476 m. From west to east, the main components of the Cretaceous to Miocene Andean fold-and-thrust belt can be observed in a regional view from this stop. To the west is a large, north-south– extending ridge known as the Cordillera del Viento. This is a basement-related structure (Fig. 17) that is mainly composed of the volcaniclastic Choiyoi Group and underlying Carboniferous metasedimentary rocks. Immediately to the east of the Cordillera del Viento is a narrow deformed belt known as the Chos Malal fold belt. Toward the east, the Chos Malal fold belt terminates against a series of large basement structures known as the Tromen synclinorium and La Yesera and Pampa Tril anticlines. These folds are characterized by flat crests and steeply dipping flanks. The insertion of the basement structures within the sedimentary sequence has produced complex triangle zones and elongated deformed fold belts.

Location: East of Laguna Auquilco at 37°23′07″S; 69°55′19″W; 1235 m. The Huantraico syncline (Fig. 11) north of the Cortaderas Lineament constitutes the eastern limit of the Agrio fold-andthrust belt. From this stop, it is possible to observe the early to middle Miocene lavas overlying the Mesozoic–lower Tertiary sedimentary sequences that fill the northern and central part of the Neuquén Basin (Fig. 18). These lavas erupted in a backarc position (Kay and Copeland, 2006) relative to the ca. 24–20 Ma Cura Mallín Formation volcanic and sedimentary sequences that formed in an intra-arc basin (e.g., Suárez and Emperán, 1995; Jordan et al., 2001; Burns, 2002; Burns et al., 2006). The backarc lavas erupted north of or along the Cortaderas lineament (Ramos, 1978), which generally marks the southern extent of post–Oligocene backarc magmatic activity in the Neuquén Basin. The significance of the Cortaderas boundary, which lies along a persistent northeast-trending regional structural grain and projects into an offset in the modern Southern Volcanic Zone arc has been unclear. Kay et al. (2006) argue that this lineament marks the southern boundary of a transient Miocene shallow subduction zone.

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Stop 1: Panoramic View of the Cordillera del Viento, Chos Malal Fold Belt, and the Tromen Volcano (North of Cortaderas Lineament)

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style is dominated by thin-skinned structures reactivated by basement inversion. The main evidence for Miocene deformation in this region comes from synorogenic deposits on the backlimb of the Pichi Mula anticline, south of the Cortaderas lineament, and from deformed lower Miocene igneous rocks in the Huantraico syncline, north of the Cortaderas lineament.

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Pliocene earliest middle to early Miocene

Scale - 10km

Parva Negra Basalt Villegas Basalt Pichi Tril Andesite Carrere Tuffs (not shown) Desfiladero Negro dikes Huantraico & Filo Morado secuences C° Cabras Basalt

Figure 18. Compilation map of the Sierra de Huantraico region from Ramos and Barbieri (1989) as modified by Kay and Copeland (2006). Map shows distribution of volcanic units, radiometric ages, and principal faults and fold axis. Open circles indicate 40Ar/39Ar ages in Kay and Copeland (2006).

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Stop 3: Cerro Rayoso Anticline

Stop 4: Puesto Burgos and Rincón Bayo Formations

Location: Two stops along Provincial Route 9 as indicated below.

Location: 38°02′6″S; 69°58′12″W; 1173 m. From this stop, a white and red stratigraphic sequence is seen near the Neuquén River (Fig. 21). These deposits are located in the frontal part of the fold-and-thrust belt in the Pampa de Agua Amarga. The red and white unit is composed of primary and reworked pyroclastic deposits and tuffs (Leanza and Hugo, 2001) that are gently folded. They unconformably overlie the Neuquén Group with an onlap relationship (Fig. 21). Zapata et al. (2002) argued that these synorogenic deposits are of middle to upper Miocene age on the basis of mammal fossils. These deposits are onlapped by the Rincón Bayo Formation, which marks the end of the deformation (Zamora Valcarce et al., 2006).

Stop 3a: Panoramic View at 37°38′11″S; 69°58′41″W; 844 m. Cerro Rayoso (Fig. 11) is a doubly plunging anticline that has an elongated, north-south–trending axis extending for more than 22 km (Fig. 19). The maximum fold width is 6 km. The lower units of the Agrio Formation crop out in the core of the anticline. There is a dramatic increase in dip from the crest toward the flanks of the anticline, with flank dips reaching 70°. The hinges are narrow and show a chevron-like structure. According to Cristallini and Allmendinger (2000), the structure is quite symmetrical based on the distribution of strain indicators. These indicators show evidence of layer-parallel shear toward the crest of the structure (a consequence of the parallel folding process). The crest is perpendicular and oblique to the actual bedding. The proposed kinematic model that explains the geometry of the structure is attributed to either detachment folding or to a trishear folding mechanism, with the former being the more probable. Stop 3b: Panoramic View at 37°38′11″S; 69°58′41″W; 844 m. At the western flank of the Cerro Rayoso anticline is the broad Pichi Neuquén syncline, which affects the uppermost units of the Rayoso Formation and the overlying Neuquén Group deposits (Fig. 20A). The syncline overlies the southern termination of the Loma Rayoso anticline. The contact between the two structures is marked by a regional fault. This fault zone, which represents the upper detachment of the Agrio fold-and-thrust belt structures, is located in the Rayoso Formation evaporites. The detachment level is so efficient that it also detaches the Loma Rayoso anticline from the Pichi Neuquén surface syncline at depth. The surface geometry relationship is imaged on the seismic lines in Figure 20B.

Figure 19. Panoramic view to the south of the Cerro Rayoso anticline.

DAY 4—LATE MIOCENE TO RECENT DEFORMATION IN THE HIGH ANDES (37°30′S) Introduction to Day 4 On Days 1–3, we visited the Agrio fold-and-thrust belt where middle Cretaceous to late Miocene contractional deformation is responsible for the observed structures (Fig. 22). Today, a complex deformational history, which began with the collapse of the inner sectors of the Agrio fold-and-thrust belt during the late Oligocene–early Miocene to form the intra-arc Cura Mallín basin (Folguera et al., 2006b), will be observed. The Oligocene to Miocene outcrops displaying this deformation are located west of the Agrio fold-and-thrust belt. The westernmost sector of the Agrio deformed belt lies beneath thick piles of volcanic rocks in the modern arc and westernmost retroarc. The main inversion of the late Oligocene–early Miocene Cura Mallín basin and its final incorporation into a fold-and-thrust belt occurred in the late Miocene (Folguera et al., 2006b), synchronous with the stacking of the external part of the Agrio fold-and-thrust belt seen on Day 3. During the final closure of the Cura Mallín basin, the Paleogene and Neogene fill was stacked to the east over the previously denuded western sector of the Agrio fold-and-thrust belt. The contraction in the inner sectors of the cordillera is considered to be related to a separate belt called the Guañacos foldand-thrust belt that is west of the Agrio belt (Figs. 22–24). The Guañacos belt shows some important differences from the Agrio belt. First, even though the Guañacos belt started to deform during the last phases of contraction to the east in the Neuquén Basin, its evolution continued into the Late Pliocene and Quaternary. By this time, compressive deformation to the east was completely fossilized. Second, the deformation in the Guañacos belt only involves Tertiary rocks at the surface. Third, shortening in the Guañacos belt is accommodated by two mechanisms: (1) inversion of late Oligocene to early Miocene extensional structures, and (2) thrusting of the sag facies of the Cura Mallín basin. The inversion of Late Triassic structures at depth that are mechanically linked to younger rift basins could also be important; this is the main mechanism of deformation in the Agrio fold-and-thrust belt.

Figure 20. (A) Photo of the regional upper detachment that detached the uppermost sequences. (B) Seismic line through the Cerro Rayoso anticline showing the structural disruption between the sequences above and below the evaporites of the Huitrín and Rayoso Formations.

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Zapata et al.

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After an initial phase of deformation in the Guañacos foldand-thrust belt, a renewed period of extensional collapse was partially superimposed on the inverted Cura Mallín basin. This collapse produced a narrow, almost 200-km–long half-graben system called the Loncopué trough. The trough formed in an extensional phase that began in the early Pliocene with reactivations occurring into the early Quaternary. To the south of the Day 4 traverse, the Plio-Quaternary normal fault bounded trough accommodates nearly 1000 m of intra-arc sequences superimposed on the main depocenter of the Cura Mallín basin. Extension associated with the Loncopué trough in this part of the Andes occurred between 5 and 1 Ma. At the latitude of the Day 4 traverse, the Loncopué lowland is of lesser magnitude, consisting of a narrow zone of Plio-Quaternary lava flows whose extensional control has been largely modified by Late Pliocene to Quaternary deformation. This late episode of contraction is now inverting the northern section of the Loncopué trough. As a result, the Guañacos foldand-thrust belt has started an eastward progression into the foreland cannibalizing the synrift sequences that were previously

in the Loncopué trough. While the Loncopué extensional stage seems to be aborted at these latitudes (37°–37°30′S), farther south (38°–39°S), neotectonics is governed by extension related to transtensional deformation. In this southern region, there are still indications of ongoing extension and related lower crustal attenuation (Yuan et al., 2006; Folguera et al., 2007). Drive across the Cordillera del Viento to Andacollo The Cordillera del Viento is part of the innermost sector of the Agrio fold-and-thrust belt north of 37°30′S. Along with the Domuyo and Tromen massifs, the Cordillera del Viento is one of the highest regions in the province of Neuquén. The range had an initial and major phase of uplift ca. 69 Ma (Burns, 2002; Burns et al., 2006) followed by minor reactivation in the late Miocene. Higher degrees of shortening in comparison with the inner part of the Agrio fold-and-thrust belt to the south (Day 2) have exposed the basement of the Neuquén Basin in this region. This metamorphic basement, which is of unknown age (Zappettini et al.,

Andean Cordillera and backarc of the south-central Andes

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Figure 22. Map of the Guañacos fold-and-thrust belt (FTB) seen on Day 4 showing the spatial relation to the Agrio fold-and-thrust belt discussed on Days 1–3 (modified from Folguera et al., 2006a). Vn—volcano.

1986), is unconformably covered by fossil-bearing Carboniferous marine sequences, which are in turn unconformably covered by the extensive rhyolites, andesites, and volcanic breccias that form the Choiyoi Group. The structure of the Cordillera del Viento has been traditionally interpreted as a major basement structure, although the style of deformation has been interpreted differently through time. Initially, this basement cored structure was thought to be connected with a system of ramps and flats that transported the basement horizontally. These structures were thought to involve large amounts of shortening and to have produced deformation in the Jurassic to Cretaceous sequences of the Neuquén Basin to the east (Kozlowski et al., 1996). More recent models relate the uplift to tectonic inver-

sion of normal faults associated with Late Triassic–Early Jurassic depocenters (Fig. 17) (Zapata et al., 1999, 2002). These faults would transport less shortening into the fold-and-thrust belt to the east than the faults considered in previous models. In detail, the Cordillera del Viento is now considered to be a back thrust, associated with a décollement that resulted from the inversion of a Late Triassic detachment (Fig. 17). Older, deeper sequences are exposed on the western slope, whereas synrift sequences and sag facies of the Neuquén Basin are seen on the eastern side. Early Jurassic sequences that form part of the synrift association (Vergani et al., 1995) are commonly grouped into the Cuyo Group and locally described as the Chacay Melehue Formation (Llambías et al., 1978). They are exposed along the southern tip of this range.

Zapata et al.

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Figure 23. Image showing locations of stops on Day 4 relative to the Guañacos fold-and-thrust belt.

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Early Pliocene to Early Quaternary extension in the Loncopué trough

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Inverted easternmost edge of the Paleogene basin Main pulse of inversion in the Cura Mallín Basin

Avalanches associated with neotectonics

EL MONCOL

THE OROGENIC FRONT

Figure 24. Enlarged view of a portion of Figure 23 showing specific features to be seen at Stops 3–5.

Stop 1: Late Triassic to Early Jurassic Depocenters of the Inner Agrio Fold-and-Thrust Belt Location: South of Cordillera del Viento at 37°17′08″S; 70°33′21″W; 1807 m. This stop is located at the southern extreme of the Cordillera del Viento, north of the Neuquén canyon. From this stop, we can see almost the complete section that has been denuded in the core of the Cordillera del Viento since the emplacement of the basement back thrust. The sequence shows the Upper Permian

to Triassic volcaniclastic Choiyoi Group covered by the Lower Jurassic marine sediments of the Cuyo Group and the Middle Jurassic to Lower Cretaceous regressive and transgressive cycles of the Neuquén embayment. The Upper Triassic to Lower Jurassic sequences were deposited during the synrift stages of a northsouth–elongated rift system that developed in the western part of the Neuquén embayment during the initial stages of its evolution. At Stop 1, unusual Lower Jurassic strata will be seen that have been interpreted as the product of a laharic mass wasting phenomena that advanced through coastal environments (Llambías

Andean Cordillera and backarc of the south-central Andes et al., 1978). Evidence for gravitational deformation that affected breccias, pyroclastic flows, and shore clastic wedges on steep slopes at the time of intraplate volcanism during the first phases of Gondwana breakup will also be seen. Drive across the Loncopué Trough Few examples of retroarc extensional basins are well preserved in the southern central and Patagonian Andes. The less than 5 Ma extensional stage of the western sectors of the Andean fold-and-thrust belt at these latitudes is represented by the Loncopué trough, which has been being mildly inverted in the past 2 m.y. in its northern section (Figs. 25 and 26). Little information exists as to the amount of collapse on the Cretaceous to Miocene structures, although Lower Jurassic strata on the western side of the Andes can be interpreted to indicate that Jurassic strata are present beneath the Loncopué trough. Extensional faults with young morphological expression border the eastern flank of the trough (García Morabito and Folguera, 2005). These faults cause eastward tilting of the early Pliocene synrift sequences that in turn produce a Quaternary half graben, which has been partially filled by a synrift depocenter. Other minor extensional and transtensional depocenters to the west, superimposed over the Guañacos fold-and-thrust belt, indicate that extension affected the whole western sector of the cordillera. Stop 2: A Quaternary Extensional Basin Developed in the Retroarc Zone between 37°10′ and 37°30′S Location: West of the town of Andacollo at 37°13′35″S; 70°42′50″W; 1074 m. At Stop 2, the upper section of the Loncopué trough fill consists of well-exposed, early Quaternary andesitic lavas. The lower Pliocene section of the basin is missing at this point but is present in depocenters to the west and south where the complete section is represented. The flat geometry of the sequences here contrasts with the highly deformed older sequences to the east and west in the Agrio and Guañacos fold-and-thrust belts. The andesitic lavas here were emplaced in a depression between two deformational belts characterized by relatively higher relief. Drive South Crossing the Guañacos River and Discussion of the Regional Picture The Cura Mallín Basin in the Guañacos River The late Oligocene to early Miocene Cura Mallín basin is characterized at these latitudes by an important asymmetry that is reflected in the sedimentation pattern. To the west, both along the western slope of the Andes and near the continental drainage divide, contractional deformation has exhumed synrift packages. These packages are identified as synextensional because the main thicknesses are associated with reverse faults with normal relationships at the surface indicating that they are

47

inverted normal faults. The sediment packages are composed of volcaniclastic materials and minor fluvial sequences. The easternmost part of the basin is superficially formed by sag sequences composed of lacustrine sediments and low-energy fluvial systems forming uniform blankets, cropping out in the upper Lileo fluvial basin (Figs. 27 and 28). These sequences were detached during the main inversion phase of the Cura Mallín basin that produced the thin-skinned deformation in the easternmost part of the Guañacos fold-and-thrust belt. Those sequences were dated as early Miocene by Sarris (1964) with palynomorphs, and related to a lacustrine environment coeval with volcanism in the Cura Mallín basin. Their lateral continuity as well as facies homogeneity makes them a good marker throughout the basin. The Contractional Orogenic Front of the Andes at 37°30′S Middle to late Miocene deformation associated with inversion of Paleogene extensional structures formed the easternmost part of the Guañacos fold-and-thrust belt. After 2–1.5 Ma, the Guañacos fold-and-thrust belt suffered an in-sequence reactivation with respect to the late Miocene deformational phase. The main activity of this last phase of deformation is considered to be early Quaternary as 1.7 Ma lavas are deformed and 1.4 Ma lavas unconformably overlie the structure. Scarps developed on soils as well as progressive unconformities in Late Quaternary sediments suggest that minor deformation on these structures persists. Only the eastern 30 km of the Guañacos fold-and-thrust belt displays evidence for Quaternary shortening. In this region, the faults show different slip sense depending on their orientation. A northeast-trending set mainly absorbed right lateral components imposed by the oblique convergence between South America and Nazca plate. North- and northwest-trending structures show reverse offsets with subordinate strike-slip displacements. These faults (Antiñir-Copahue fault system or the orogenic front of the Guañacos fold-and-thrust belt) show a curved geometry in plan view, which is delimited by two regional lineaments associated with the Neuquén and Cura Mallín basins (Fig. 29). These are the Mandolegüe lineament to the south and the Cortaderas lineament to the north. Stop 3: Folding and Thrusting in Lower Quaternary Retroarc Volcanics and Upper Quaternary Sediments: The Antiñir-Copahue Fault System at the Orogenic Front of the Guañacos Fold and Thrust and the Main Phase of Inversion of the Cura Mallín Basin Location: Reñileuvú River at 37°20′39″S; 70°48′52″W; 1177 m. At this stop on the Reñileuvú River, one of the clearest expressions of the orogenic front at the latitude of 37°30′S is visible. Here, 1.7 Ma lavas are folded in an anticline (Fig. 30). This anticline is offset by a reverse fault in the section seen along the Guañacos River section to the north. An east-dipping monocline formed at the structural front affects Upper Quaternary

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Figure 25. (A) View of the Guañacos River at Stop 2 on Day 4. (B) Map of the Loncopué trough in the retroarc area of the Andes between 36° and 39°S from Ramos and Folguera (2005).

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Figure 26. Geological map of the section of the Loncopué trough seen along the transect of Day 4 of this field trip. Map shows relation of the Loncopué trough to the compressive structures along the orogenic front (modified from Folguera et al., 2004).

Synorogenic Quaternary deposits

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Clastic sedimentary rocks of the Cura Mallín Formation Breccias and lavas of the Cura Mallín Formation Tuffs of the Cura Mallín Formation

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Figure 27. Structural cross section along the Lileo and Guañacos rivers showing the contractionally inverted structures associated with Miocene synrift wedges related to half grabens and sag sequences in the easternmost part of the Guañacos fold-and-thrust belt (modified from Folguera et al., 2006a).

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Figure 28. Geologic map showing a section of the eastern slope of the Andes between the Reñileuvú and Guañacos valleys from Folguera et al., 2006a). Vn—volcano.

sediments, which are associated with an outwash fan related to Pleistocene glaciation. In the frontal syncline, a narrow tongue of synorogenic deposits thickens to the east into the fold axis. Late Oligocene collapse of the inner sectors of the foldand-thrust belt at these latitudes was followed by mild inversion between 20 and 17 Ma, and final closure of the intra-arc basin at 10–8 Ma. This deformation can be observed in the inner western sectors of the cordillera on a regional unconformity between synrift sequences and early Pliocene plateau successions. In contrast, the contractional deformation that reinitiated in the Quaternary along the easternmost sector of the Guañacos fold-and-thrust belt exhibited at Stop 3 has progressed to the east with no substantial deformation to the west. This eastward progression explains why lower Pliocene sequences in the high cordillera associated with the extensional stages of the Loncopué trough have not been contractionally deformed.

Stop 4: Main Unconformity between Lower Miocene Synrift Sequences and Pliocene Units Location: Along the Reñileuvú River at 37°20′35″S; 70°52′50″W; 1223 m. Stop 4 is located in the westernmost Argentinean side of the Andes at 37°30′S. At this stop, synrift sequences of the Cura Mallín basin, which are mainly composed of volcanic breccias, lacustrine sediments, and subordinate lava flows, are unconformably

Figure 29. Schematic geological map of the Antiñir-Copahue fault system from Folguera et al. (2006a). Map shows the orogenic front of the Guañacos fold-and-thrust belt that resulted from the inversion of the Cura Mallín basin. Rectangle in the center of the map shows the location of the transect along the Reñileuvú River on Day 4 relative to the map and image in Figures 24 and 28. Vn—volcano.

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Figure 30. Photograph and sketch of the orogenic front of the Guañacos fold-and-thrust belt as seen from Stop 4 on Day 4.

overlain by Early Pliocene plateau lavas (Fig. 31). Nearly flatlying Pliocene sequences show that contraction in this area was over before the Early Pliocene. Stop 5: Mass Wasting Phenomena Associated with the Orogenic Front and Geological Hazards Associated with Seismic-Triggered Avalanches in the Moncol Area Location: Nera Moncol at 37°21′42″S; 71°00′06″W; 1467 m. An effect of the young tectonic activity throughout this region is the alignment of avalanche deposits along traces of neotectonic faults. The amount of mass wasting is anomalous compared to other parts of the Andean fold-and-thrust belt due to the position of the Guañacos fold-and-thrust belt. The migration of the orogenic front toward the west during the latest Pliocene rather than toward the foreland as in most of the rest of the Andes has

produced young faults. These faults are intercepting the deeply carved, glacial-fluvial basin and reactivating a few of the late Miocene faults such as the El Moncol fault in the region of this stop (Figs. 28 and 32). This situation has produced steep slopes on the Pleistocene glacial morphology that have become gravitationally destabilized by both erosion and crustal earthquakes. The innermost neotectonic faults in the Moncol area are associated with profuse mass wasting phenomena (Fig. 32). Some associated features of the cordillera in this region need to be highlighted: (1) a large magnitude of mass wasting that has profoundly altered the valley morphology; (2) a minor post-glacial fluvial incision that contrasts with sectors to the east where fluvial processes have carved a glacial morphology; and (3) changes in valley morphology that are spatially related to neotectonic faults along the orogenic front of the Guañacos fold-and-thrust belt. Drive Back to Chos Malal through the Cortaderas Lineament (Neuquén Canyon)

Early Pliocene lavas

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The stops on the last part of Day 4 on the way back to Chos Malal are again in the westernmost Agrio fold-and-thrust belt. Along this part of the route, the Middle to Upper Jurassic continental sequences that extensively cover the synrift deposits of the Neuquén Basin are well exposed. After a short period of evaporation during the Middle Jurassic when extensive salty environments developed in relation to a fall in sea level, fluvial systems dominated the Neuquén Basin. Stop 6: Late Jurassic Depocenter at the Inner Part of the Agrio Fold-and-Thrust Belt

Río Reñileuvú

Figure 31. Regional unconformity between the early Pliocene plateau lava sequences related to the Loncopué extensional stage and late Oligocene to early Miocene sedimentary beds, which accumulated as synrift wedges and were contractionally inverted in the late Miocene.

Location: Neuquén River to Chos Malal at 37°22′44″S; 70°30′38″W; 1167 m. Passing through El Mollar along the southern margin of the Neuquén River, the southernmost extreme of the Cordillera del Viento is crossed south of Stop 1 on Day 4. Here more than 1000 m of Late Jurassic continental Tordillo Group sandstones cover synrift sequences of the Neuquén Basin (Fig. 33). They are

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Figure 32. Principal avalanches seen along the Reñileuvú River and Guañacos valley. Note that mass wasting phenomena are present in broadened glacial valleys, whereas neotectonic activity is associated with deeply incised canyons carved after Pleistocene glaciations.

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Main mass wasting phenomena through the region A. Break away zone B. Direction of avalanche emplacement C. Displaced rock mass

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Figure 33. Three-dimensional Landsat image from Google Earth through the Neuquén Canyon, along whose trace, anomalously thick sequences of continental sandstones in the Tordillo Formation are found.

anomalously thick as can be achieved only in a basin. These Late Jurassic continental sandstones represent a low stand of sea level that is seen throughout the basin. The sands accumulated on erosional surfaces like those in deep channels. The Cortaderas lineament that runs through the southern extreme of the Cordillera del Viento is thought to have had a strong influence on differential fluvial incision. END OF TRIP.

REFERENCES CITED Bettini, F.H., 1982, Complejos efusivos terciarios presentes en las Hojas 30c y 32b (Puntilla de Huincan y Chos Malal, del sur de Mendoza y norte de Neuquén), Argentina: V Congreso Latinoamericano Geológico, Argentina, Actas, v. 5, p. 79–114. Bodenbender, G., 1891, Apuntes sobre rocas eruptivas de la pendiente oriental de los Andes entre el Río Diamante y Río Negro: Revista Argentina de Historia Natural, Tomo I.

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Burckhardt, C., 1900, Coupe geólogiques de la Cordillere entre las Lajas et Curacautín: Anales Museo La Plata, Sección Mineralogía y Geología, v. 3, p. 1–102. Burns, W.M., 2002, Tectonics of the southern Andes from stratigraphic, thermochronologic, and geochemical perspectives [Ph.D. thesis]: Ithaca, New York, Cornell University, 204 p. Burns, W.M., and Jordan, T.E., 1999, Extension in the southern Andes as evidenced by an Oligo-Miocene age intra-arc basin, in Andean geodynamics: Fourth International Symposium on Andean Geodynamics meeting, Gottingen, Germany: Paris, Institut de Recherche pour le Developpement, p. 115–118. Burns, W.M., Jordan, T.E., Copeland, P., and Kelley, S.A., 2006, The case for extensional tectonics in the Oligocene-Miocene southern Andes as recorded in Cura Mallín basin (36°–38°S), in Kay S.M., and Ramos V.A., eds., Evolution of an Andean margin: A tectonic and magmatic view from the Andes to the Neuquén Basin (35°–39°S lat): Geological Society of America Special Paper 407, p. 163–184. Cobbold, P.R., and Rosello, E.A., 2003, Aptian to recent compressional deformation, foothills of the Neuquén Basin Argentina: Marine and Petroleum Geology, v. 20, p. 429–443, doi: 10.1016/S0264-8172(03)00077-1. Cristallini, E.O., and Allmendinger, R.W., 2000, Estructura de la Faja plegada del Agrio, Provincia del Neuquén: Repsol YPF internal report, 68 p. Folguera, A., Ramos, V., Hermanns, R., and Naranjo, J., 2004, Neotectonics in the foothills of the southernmost central Andes (37°–38°S): Evidence of the strike-slip displacement along the Antiñir-Copahue fault zone: Tectonics, v. 23 (TC 5008), 23 p. Folguera, A., Ramos, V.A., Gonzalez Díaz, E.F., and Hermanns, R., 2006a, Miocene to Quaternary deformation of the Guañacos fold-and-thrust belt in the Neuquén Andes between 37°S and 37°30′S, in Kay, S.M., and Ramos, V.A., eds., Evolution of an Andean margin: A tectonic and magmatic view from the Andes to the Neuquén Basin (35°–39° lat): Geological Society of America Special Paper 407, p. 247–266, doi: 10.1130/2006.2407(11). Folguera, A., Zapata, T., and Ramos, V., 2006b, Late Cenozoic evolution of the eastern Andean foothills of Neuquén, in Kay, S.M, and Ramos, V.A., eds., Evolution of an Andean margin: A tectonic and magmatic view from the Andes to the Neuquén Basin (35°–39°S lat): Geological Society of America Special Paper 407, p. 267–285. Folguera, A., Introcaso, A., Giménez, M., Ruiz, F., Martínez, P., Tunstall, C., García Morabito, E., and Ramos, V.A., 2007, Crustal attenuation in the southern Andean retroarc determined from gravimetric studies (38°–39°30′S): The Lonco-Luán asthenospheric anomaly: Tectonophysics: v. 439, p. 129–147, doi: 10.1016/j.tecto.2007.04.001. García Morabito, E., and Folguera, A., 2005, El alto de Copahue-Pino Hachado y la fosa de Loncopué: Un comportamiento tectónico episódico, Andes neuquinos (37°–39°S): Asociación Geológica Argentina, Revista, v. 60, no. 4, p. 742–761. Gerth, E., 1928, Estructura geológica de la Cordillera Argentina entre los ríos Grande y Diamante en el sur de la provincia de Mendoza: Córdoba, Actas, Academia Nacional de Ciencias, v. 10, p. 122–170. Groeber, P., 1929, Líneas fundamentales de la Geología de Neuquén, sur de Mendoza y regiones adyacentes: Buenos Aires, Ministerio de Agricultura, Dirección General de Minas, Geología e Hidrología Publicación, v. 58, p. 1–10. Groeber, P., 1946a, Observaciones geológicas a lo largo del meridiano 70. I: Hoja Chos Malal: Sociedad Geológica Argentina, Revista, v. 1, no. 2, p. 177–208. Groeber, P., 1946b, Observaciones geológicas a lo largo del meridiano 70. II: Hojas Domuyo, Mari-Mahuida, Huarhuar Co y parte de Epu Lauken: Sociedad Geológica Argentina, Revista, v. 2, no. 4, p. 347–408. Gulisano, C.A., Gutierrez Pleimling, A.R., and Digregorio, J.H., 1984, Análisis estratigráfico del intervalo Tithoniano-Valanginiano (formaciones Vaca Muerta, Quintuco y Mulichinco) en el suroeste de la Provincia del Neuquén: San Carlos de Bariloche, IX Congreso Geológico Argentino, Actas, v. 1, p. 221–235. Gutiérrez, P.A., and Minniti, S., 1985, Reconocimiento geológico de las nacientes del Río Lileo (departamento Minas), provincia del Neuquén, Argentina: Buenos Aires, YPF, unpublished report. Haller, M., Nullo, F.E., Proserpio, C.A., Parica, P.D., and Cagnoni, M.C., 1985, Major element geochemistry of early Tertiary volcanics: Santiago, Departamento de Geología, Universidad de Chile, Comunicaciones, v. 35, p. 97–100.

Herrero Ducloux, A., 1946, Contribución al conocimiento geológico del Neuquén extrandino: Boletín de Informaciones Petroleras, v. 23, no. 226, p. 1–39. Jordan, T.E., Burns, W.M., Veiga, R., Pángaro, F., Copeland, P., Kelley, S., and Mpodozis, C., 2001, Extension and basin formation in the southern Andes caused by increased convergence rate: A mid-Cenozoic trigger for the Andes: Tectonics, v. 20, p. 308–324, doi: 10.1029/1999TC001181. Kay, S.M., and Copeland, P., 2006, Early to middle Miocene backarc magmas of the Neuquén Basin: Geochemical consequences of slab shallowing and the westward drift of South America, in Kay, S.M., and Ramos, V.A., eds., Evolution of an Andean margin: A tectonic and magmatic view from the Andes to the Neuquén Basin (35°–39°S lat): Geological Society of America Special Paper 407, p. 185–213. Kay, S.M., Burns, W.M., Copeland, P., and Mancilla, O., 2006, Upper Cretaceous to Holocene magmatism and evidence for transient Miocene shallowing of the Andean subduction zone under the northern Neuquén Basin, in Kay, S.M., and Ramos, V.A., eds., Evolution of an Andean margin: A tectonic and magmatic view from the Andes to the Neuquén Basin (35°–39°S lat): Geological Society of America Special Paper 407, p. 19–60. Kozlowski, E., Cruz, C.E., and Rebay, G.A., 1987, El Terciario vulcanoclástico de la zona de Puntilla de Huincán, provincia de Mendoza, Argentina: San Miguel de Tucumán, X Congreso Geológico Argentino, Actas, v. IV, p. 229–232. Kozlowski, E., Cruz, C.E., and Sylwan, C.A., 1996, Geología estructural de la zona de Chos Malal: Cuenca Neuquina, Argentina: Buenos Aires, XIII Congreso Geológico Argentino y III Congreso de Exploración de Hidrocarburos, Actas, v. I, p. 15–26. Leanza, H.A., and Hugo, C.A., 2001, Hoja geológica 3969-I: Zapala, provincia del Neuquén: Buenos Aires, Programa Nacional de la Carta Geológica, Servicio Geológico y Minero Argentino (SEGEMAR), escala 1:100,000. Legarreta, L., and Gulisano, C.A., 1989, Análisis estratigráfico secuencial de la cuenca neuquina (Triásico Superior-Terciario Inferior), Argentina, in Chebli, G.A., and Spalleti, L.A., eds., Cuencas sedimentarias Argentinas: Serie Correlación Geológica, v. 6, p. 21–243. Llambías, E.J., and Malvicini, L., 1978, Geología, petrología y metalogénesis del área de Colipilli, provincia del Neuquén, República Argentina: Asociación Geológica Argentina, Revista, v. 33, p. 257–276. Llambías, E.J., and Rapela, C.W., 1987, Las vulcanitas de Colipilli y sus relaciones con las provincias volcánicas del Terciario inferior de NeuquénMendoza y Patagonia: San Miguel de Tucumán, X Congreso Geológico Argentino, Actas, v. 4, p. 249–251. Llambías, E.J., and Rapela, C.W., 1989, Las vulcanitas de Colipilli, Neuquén (37°S) y su relación con otras unidades paleógenas de la cordillera: Asociación Geológica Argentina, Revista, v. 44, no. 1–4, p. 224–236. Llambías, E.J., Danderfer, J.C., Palacios, M., and Brogioni, N., 1978, Las rocas ígneas cenozoicas del Volcan Domuyo y áreas adyacentes: VII Congreso Geológico Argentino (Neuquén), Actas, v. 2, p. 569–584. Niemeyer, H., and Muñoz, J., 1983, Geología de la hoja 57 Laguna de la Laja, Región de Bío Bío: Santiago, Chile, Servicio Nacional de Geología y Minería, scale 1:250,000, 1 sheet. Pesce, A., 1981, Estratigrafía de las nacientes del río Neuquén y Nahuever Provincia del Neuquén: San Luis, VIII Congreso Geológico Argentino, Actas, v. 3, p. 439–455. Ramos, V.A., 1977, Estructura, de la Provincia de Neuquén, in Rolleri, E.O., ed., Geología y recursos naturales de la Provincia del Neuquén: Buenos Aires, VII Congreso Geológico Argentino (Neuquén), Asociación Geológica Argentina, Relatorio de la Provincia de Neuquén, p. 9–24. Ramos, V.A., 1978, Estructura, in Relatorio de la geología y recursos naturales del Neuquén: Buenos Aires, VII Congreso Geológico Argentino, p. 99–118. Ramos, V.A., 1981, Descripción geológica de la Hoja 33c, Los Chihuidos Norte: Buenos Aires, Boletín del Servicio Geológico Nacional, v. 182, p. 1–103. Ramos, V.A., 1989, The birth of southern South America: American Scientist, v. 77, p. 444–450. Ramos, V.A., 1998, Estructura del sector occidental de la faja plegada y corrida del Agrio, cuenca Neuquina, Argentina: Buenos Aires, X Congreso Latinoamericano de Geología, Actas, v. 2, p. 105–110. Ramos, V.A., 1999, Rasgos estructurales del Territorio Argentino, Evolucion Tectonica de la Argentina: Buenos Aires, Geologia Argentina, Instituto de Geologia y Recursos Naturales, Anales, v. 29, p. 715–784.

Andean Cordillera and backarc of the south-central Andes Ramos, V.A., and Barbieri, M., 1989, El volcanismo Cenozoico de Huantraico: Edad y relaciones isotópicas iniciales, provincia de Neuquén: Buenos Aires, Revista Asociación Geológica Argentina, v. 43, p. 210–223. Ramos, V.A., and Folguera, A., 2005, Tectonic evolution of the Andes of Neuquén: Constraints derived from the magmatic arc and foreland deformation, in Veiga, G., Spalletti, L., Howell, J., and Schwarz, E., eds., The Neuquén Basin: A case study in sequence stratigraphy and basin dynamics: Geological Society [London] Special Publication 252, p. 15–35. Ramos, V.A., and Kay, S.M., 1992, Southern Patagonian plateau basalts and deformation: backarc testimony of ridge collisions: Tectonophysics, v. 205, p. 261–282, doi: 10.1016/0040-1951(92)90430-E. Rapela, C.W., and Llambías, E.J., 1985, La secuencia andesítica Terciaria de Andacollo, Neuquén, Argentina: IV Congreso Geológico Chileno (Antofagasta), Actas, v. 3, p. 4-458–4-488. Repol, D., Leanza, H.A., Sruoga, P., and Hugo, C.A., 2002, Evolución tectónica del Cenozoico de la comarca de Chorriaca, Provincia del Neuquén, Argentina, in Cabaleri, N., Cingolani, C.A., Linares, E., López de Luchi, M.G., Ostera, H.A, and Panarello, H.O., eds.: XV Congreso Geológico Argentino (El Calafate), Actas CD-ROM, Artículo 227, 6 p. Rovere, E.I., and Rossello, E., 2001, Evolución geológica durante el Miocene en la región del C° Columpios, 37°S, Andes neuquinos, Argentina: XI Congreso Latinoamericano (Montevideo, Uruguay), Actas CD-ROM. Sarris, M., 1964, Informe geológico de la zona de río Palao, departamento minas, provincia del Neuquén: Yacimientos Carboníferos Fiscales Open-File Report 907. Suárez, M., and Emperán, C., 1995, The stratigraphy, geochronology and paleophysiography of a Miocene fresh-water interarc basin, southern Chile: Journal of South American Earth Sciences, v. 8, no. 1, p. 17–31, doi: 10.1016/0895-9811(94)00038-4. Uliana, M., and Legarreta, L., 1993, Hydrocarbon habitat in a Triassicto-Cretaceous sub-Andean setting: Neuquén basin, Argentina: Journal of Petroleum Geology, v. 16, p. 397–420, doi: 10.1111/j.1747-5457.1993. tb00350.x. Vergani, G.D., Tankard, A.J., Belotti, H.J., and Welsink, H.J., 1995, Tectonic evolution and paleogeography of the Neuquén basin, Argentina, in Tankard, A.J, Suarez, R.S., and Welsink, H.J. eds., Petroleum basins of South America: American Association of Petroleum Geologists Memoir 62, p. 383–402. Viñes, R.F., 1985, Estilos estructurales de la faja plegada occidental neuquina: Repsol-YPF, unpublished report, 6 p.

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Weaver, C.E., 1931, Paleontology of the Jurassic and Cretaceous of west-central Argentina: Memoir of the University of Washington, v. 1, p. 1–496. Yrigoyen, M.F., 1972, Cordillera Principal, in Leanza, A., ed., Geología regional Argentina: Córdoba, Academia Nacional de Ciencias, p. 345–364. Yuan, X., Asch, G., Bataile, K., Bohm, M., Echtler, H., Kind, R., Onchen, O., and Wölbern, I., 2006, Deep seismic images of the southern Andes, in Kay, S.M., and Ramos, V.A., eds., Evolution of an Andean margin: A tectonic and magmatic view from the Andes to the Neuquén Basin (35°–39°S lat): Geological Society of America Special Paper 407, p. 61–72. Zamora Valcarce, G., 2007, Estructura y cinemática de la faja plegada del Agrio [Ph.D. thesis]: Buenos Aires, University of Buenos Aires, 302 p. Zamora Valcarce, G., Zapata, T.R., and del Pino, D., and Ansa, A., 2006, Structural evolution and magmatic characteristics of the Agrio fold-andthrust belt, in Kay, S.M., and Ramos, V.A., eds., Evolution of an Andean margin: A tectonic and magmatic view from the Andes to the Neuquén Basin (35°–39°S lat): Geological Society of America Special Paper 407, p. 125–145. Zamora Valcarce, G., Rapalini, A., and Spagnuolo, C., 2007, Reactivación de estructuras cretácicas durante la deformación miocena: Faja plegada del Agrio, Neuquén: Revista de la Asociación Geológica Argentina, v. 62, no. 2, p. 299–308. Zanettini, J., Méndez, V., and Zappettini, E., 1987, El Mesozoico y Cenozoico sedimentario de la comarca de los Miches, Provincia de Neuquén: Revista de la Asociación Geológica Argentina, v. 42, no. 3–4, p. 338–348. Zapata, T.R., and Folguera, A., 2005, Tectonic evolution of the Andean fold and thrust belt of the southern Neuquén Basin, Argentina, in Veiga, G., Spalletti, L., Howell, J., and Schwarz, E., eds., The Neuquén Basin: A case study in sequence stratigraphy and basin dynamics: Geological Society [London] Special Publication 252, p. 37–56. Zapata, T.R., Brissón, I., and Dzelalija, F., 1999, The role of basement in the Andean fold and thrust belt of the Neuquén Basin, in McClay, K., ed., Thrust tectonics 1999: London, v. I, Abstracts, p. 122–124. Zapata, T.R., Córsico, S., Dzelalija, F., and Zamora Valcarce, G., 2002, La faja plegada y corrida del Agrio: Análisis estructural y su relación con los estratos Terciarios de la Cuenca Neuquina, Argentina: Mar del Plata, V Congreso de Exploración y Desarrollo de Hidrocarburos, CD-ROM. Zappettini, E., Méndez, V., and Zanettini, J., 1986, Metasedimentitas mesoplaleozoicas en el noroeste de la Provincia del Neuquén: Revista de la Asociación Geológica Argentina, v. 42, p. 206–207. MANUSCRIPT ACCEPTED BY THE SOCIETY 10 JANUARY 2008

Printed in the USA

The Geological Society of America Field Guide 13 2008

Field trip guide: Frontal and Main Andean Cordilleras near the southern boundary of the Pampean shallow subduction zone Víctor A. Ramos* Laboratorio de Tectónica Andina, Universidad de Buenos Aires, Consejo Nacional de Investigaciones Científicas y Técnicas (CONICET), Buenos Aires, Argentina ABSTRACT This one-day field trip provides a sampling of the main components of the Andean deformation front in the Precordillera, and the Frontal and Main Cordilleras of the central Andes east of the drainage divide, which at these latitudes coincide with the political boundary of Argentina and Chile. The absence of Pliocene to Recent volcanic rocks in this transect over the southern hinge of the modern shallow subduction allows the older Andean rocks and their structures to be well seen. The structural consequences of shallow subduction are also well seen. The field trip stops provide a view of the Late Paleozoic sedimentary and magmatic rocks of the Frontal Cordillera, the Triassic volcanic and plutonic sequences and associated sedimentary rift sequences east of the Main Cordillera, the Mesozoic to Miocene arc magmatic and sedimentary basin sequences of the high Cordillera, and the Miocene foreland basin deposits to the east. The structure of the Triassic Cuyo rift and inverted normal faults is contrasted with both the Miocene thick-skinned contractional structures affecting the dominantly magmatic rocks of the Frontal Cordillera and the thin-skinned folds and thrusts of the Aconcagua belt affecting the Jurassic to Miocene sedimentary and volcanic rocks of the Principal Cordillera. Depending on climatic conditions, Cerro Aconcagua (6967 m above sea level), the highest peak in the Western and Southern hemispheres and the top of the Backbone of the Americas, can be viewed. Keywords: Andes, tectonics, terrane, allochthonous, Aconcagua. INTRODUCTION

plete traverses of the Andes, where a noncollisional orogenic belt reaches elevations near 7 km (the highest mountains of the Western Hemisphere). These mountains are in an area of no present volcanic activity, and the Late Cenozoic shortening is directly related to the present uplift and convergence rate. The route as chosen will show the different structural styles of the Precordillera thrust front, the Cordilleras Frontal and Principal as indicated on the map in Figure 1, with some classic stops along the road. The route follows Highway 7, which is the major road crossing the Andes west of Mendoza.

This field trip provides the opportunity to examine the tectonic evolution of the central Andes in one of the most classic sections. The trip aims to show the key localities where field data have been obtained and the different regional relationships and tectonic models have been based (Fig. 1). The main objective of this one-day trip is to view one of the most com*[email protected]

Ramos, V.A., 2008, Field trip guide: Frontal and Main Andean Cordilleras near the southern boundary of the Pampean shallow subduction zone, in Kay, S.M., and Ramos, V.A., eds., Field trip guides to the Backbone of the Americas in the southern and central Andes: Ridge collision, shallow subduction, and plateau uplift: Geological Society of America Field Guide 13, p. 57–76, doi: 10.1130/2008.0013(03). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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Figure 1. Index map of field trip road from Mendoza to Las Cuevas.

Geologic understanding of the central Andes started with the pioneering explorations of Charles Darwin in 1835, who was the first to describe the marine Mesozoic deposits deformed by faults (Fig. 2) along the present road that crosses the High Andes. We are going to have the opportunity to examine the profile described by Darwin, and see the mountain shelters where Darwin stayed during his crossing. The early observations of Burmeister in 1857–1858, Stelzner in 1873, and Wherli and Burckhardt in 1898, presented the concept that the High Andes of San Juan and Mendoza was a relatively simple mountain chain without the thrusts and overthrusts, known in other mountain chains at the time. The work in 1906 and 1907 of Walter Schiller, a young geologist and mountain climber in the Argentine Geological Survey, revealed many structural complexities and important thrusts and changed this concept. At the same time, the Precordillera was being explored, making this region one of the best known at the time. Over the last several decades, many studies have provided a large volume of

information on different aspects of the sedimentology, structure, and geologic evolution of this part of the central Andes (see review in Caminos, 2000). See chapter 4 of this volume for further discussion of the history of exploration and references. MAJOR GEOLOGICAL PROVINCES The central Andes at this latitude have been divided into a series of geological provinces based on structural styles, geologic evolution, and morphological expression. These provinces include the block-faulted Sierras Pampeanas (González Bonorino, 1950), the Precordillera fold-and-thrust belt, the thick-skinned and rigid Frontal Cordillera, the Main Cordillera with the thin-skinned Aconcagua fold-and-thrust belt and the pre–9 Ma volcanic arc, and the Chilean coastal cordilleras to the west (see Jordan et al., 1983a, 1983b). The Precordillera, Frontal Cordillera, and Main Cordillera, which will be visited on this trip, are briefly described below. A discussion of all of the provinces with illustrations can

Frontal and main Andean Cordilleras

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Figure 2. Structural sketch section of Puente del Inca after Darwin (1846).

be found in chapter 4 of this volume. Stops in chapter 4 feature the westernmost Sierras Pampeanas and the Precordillera on Days 1–3 and the Cordillera Frontal, Cordillera Principal, and Chilean coastal Cordillera on Days 4 and 5.

1991). The Cordillera Frontal behaved as a rigid block during the Andean deformation as shown by the presence of thickskinned thrusts (Ramos et al., 1996a, 1996b). Cordillera Principal

Precordillera The Precordillera is an Andean fold-and-thrust sequence with a typical thin-skinned structure that developed in a Paleozoic platform sequence. The western edge coincides with the north-south–trending depression known as the Iglesia-CalingastaUspallata valley. The Early Paleozoic history of the Precordillera is represented by an Early Cambrian to Middle Ordovician carbonate platform sequence that is covered by late Early Paleozoic platform facies in the eastern and central sectors and slope and oceanic facies in the west. These sequences were deformed during the major Chanic deformational event in the Middle to Late Devonian. Overlying Late Paleozoic sequences are mainly littoral marine facies in the west and continental fluvial to alluvial deposits in the central and eastern sectors. Synorogenic Tertiary deposits show that Andean folding and thrusting began in the Precordillera at ca. 18 Ma in the west and is still active in the east. Cordillera Frontal The Cordillera Frontal is made up of Late Paleozoic to early Mesozoic units. Most exposures are Late Paleozoic– Triassic andesitic to silicic magmatic rocks of the Choiyoi Group (Caminos, 1979). Volcanic activity started in the Middle Carboniferous with subduction-related andesites, dacites, and rhyolites. The Carboniferous to Early Permian units were then deformed in the middle Permian San Rafael orogenic phase (Caminos, 1979; Ramos, 1988a). Generalized extension subsequently took place from the middle Permian to the Early Triassic as the thick pile of Choiyoi Formation rhyolites and associated granites were emplaced (see Kay et al., 1989; Llambías and Sato, 1990). The volcanic sequences have thicknesses up to 2–4 km along the field trip route in the Río Mendoza valley. The boundary between the Cordillera Frontal and the Precordillera was a locus of Triassic rifting, which produced up to 2 km of synrift deposits and scattered alkaline basalts (Ramos and Kay,

The Cordillera Principal or Main Andes was the locus of Andean orogeny and volcanism from the latest Mesozoic to the late Miocene (see Mpodozis and Ramos, 1990). The style of Jurassic and Cretaceous sedimentary and volcanic units controls the style of basement deformation. Several sedimentary cycles from the Early Jurassic to the Early Cretaceous are recognized. They begin with black shales, sandstones, and limestones and terminate with thick gypsum levels and continental red beds. Abundant ammonites permit a biostratigraphic zonation of these deposits (Aguirre Urreta and Rawson, 1997). Along the border with Chile, the sedimentary sequences interfinger with important Late Jurassic–Early Cretaceous volcanic and pyroclastic sequences. The volcanic pile can be up to 6 km thick on the Chilean side where it has been affected by hydrothermal activity during active subsidence (see Levi et al., 1982). These Mesozoic deposits are involved in the Neogene thin-skinned Aconcagua fold-and-thrust belt (Yrigoyen, 1976, 1979; Ramos et al., 1996a, 1996b) that is well developed on the Argentine side. A series of frontal volcanic arcs shifted from the Cordillera de la Costa in Chile in the Jurassic to the Cordillera Principal in the Miocene. Glacial deposits from four different glaciations are widespread in the main valleys, representing an alpine-type glaciation during Pliocene and Quaternary times. NEOGENE PLATE TECTONIC SETTING The central Andes between 28° and 33°S have a distinctive plate tectonic setting. Earthquake locations in this segment delineate a modern Benioff zone that is gently dipping to the east defining a shallow subduction zone (Cahill and Isacks, 1992), flanked to the north and south by segments dipping at ~30°. The present convergence rate between the subducted Nazca plate and the South America plate averages ~9 cm/yr. An obvious and consistent correlation exists between the shallowest part of the subduction zone, the absence of Quaternary volcanism,

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and a broad zone of deformation reaching far into the foreland. Present tectonic shortening is principally concentrated along a narrow north-south–trending belt between the Sierras Pampeanas and the Precordillera. Intraplate earthquake nests occur in the basement of the eastern Precordillera and the western Sierras Pampeanas, in close coincidence with crustal neotectonic activity. Focal mechanisms indicate east-west contraction with null to minor strike-slip displacements (Chinn and Isacks, 1983; Pardo et al., 2000, 2002). The stops on this field trip are over the southern hinge of the shallow subduction zone. The shallowing of the subduction zone is considered to have begun at ca. 18 Ma with the most intense phase in the late Miocene and Pliocene after ca. 10 Ma (e.g., Kay et al., 1991). The causes of the change in Benioff zone geometry and segmentation are likely multifaceted. The breakup of the Farallon plate into the Cocos and Nazca plates at 25 Ma marks the beginning of a period of higher convergence rates and the initiation of the modern Andean orogeny (Handschumacher, 1976; Pardo Casas and Molnar, 1987). This is the time of major reinitiation of volcanism and is a milestone in the geodynamic evolution of the area. A deceleration of the plate convergence rate between the subducted Nazca plate and the South America plate began at 10 Ma (Pilger, 1984; Pardo Casas and Molnar, 1987). The main origin of the flat subduction segment and overlying features is best attributed to subduction of aseismic ridges (Pilger, 1981) and differential shortening of a previously weakened continental crust (Isacks, 1988). Pilger (1981) argued that buoyancy effects produced by subduction of ridges contributed to diminishing the angle of the Benioff zone. The effects of the subduction of the Juan Fernandez hot spot trace that began at this latitude at ca. 10 Ma correlate with a southward shift of the Juan Fernandez ridge collision along the trench (von Huene et al., 1997; Yañez et al., 2001). PALEOZOIC HISTORY The response to Andean deformation in the region is influenced by preexisting units and deformational events. The Early Paleozoic history of this region is dominated by arc magmatism along a continental margin and the events associated with the amalgamation of two principal terranes called the Precordillera and Chilenia to the Gondwana margin (e.g., Ramos, 1988b). The events are briefly described below. More detail and field trip stops that feature the evidence can be found in chapter 4 of this volume. Ordovician Accretion of the Precordillera (Cuyania) Terrane The existence of oceanic rocks separating the Cordillera Frontal from the Precordillera has attracted the attention of geologists since the early work of Borrello (1969). These oceanic rocks have been interpreted as indicating a suture between two continental terranes (Ramos et al., 1984, 1986). The first report

of Borrello in 1963 of North American olenellid trilobites in the Precordillera near San Juan led to a puzzle about connections between the Laurentia and Gondwana continents. A striking coincidence in the subsidence curves of the Appalachian and Precordillera carbonate platforms led Bond et al. (1984) to suggest that both were conjugate margins and shared a common rift drift transition. Ramos et al. (1986) proposed that the Precordillera was a far-traveled terrane derived from the northern Appalachians. Astini et al. (1995, 1996) developed the model within the tectonic framework of South America, and Thomas and Astini (e.g., 1999, 2003) and Thomas et al. (2002) made comparisons with the North America margin and argued that the Precordillera was derived from the Ouachita region. Two models were proposed to explain the accretion of the Precordillera to Gondwana. One was a microcontinent detached from Laurentia during the early Cambrian and collided with Gondwana in the middle to late Ordovician (Ramos et al., 1986; Benedetto and Astini, 1993; Benedetto et al., 1999; Astini et al., 1995, 1996). The other was that Laurentia and Gondwana collided in the early to mid Ordovician and that the two continents separated in the late Ordovician leaving the Precordillera terrane behind and opening an ocean on the western side (Dalla Salda et al., 1992a, 1992b; Dalziel, 1997; Dalziel et al., 1996). Both proposals required an active early Paleozoic margin in the Sierras Pampeanas and linked the Ordovician Ocloyic deformational event to collisional orogeny (Ramos et al., 1986; Dalla Salda et al., 1992a, 1992b). The reconstruction of the late Proterozoic continents in Moores (1991) complemented by the Laurentian end run model in Dalziel (1991) provided a context for the transfer of the Precordillera. Devonian Accretion of the Chilenia Terrane The Precordillera terrane hypothesis requires a second early Paleozoic accretion to close the ocean bounding the western Precordillera. The proposal for the accretion of a terrane called Chilenia provided an explanation for a second foreland basin and a westward shift in the arc magmatic front (Ramos et al., 1984, 1986). Sedimentologic studies of the foreland basin and ages of peak metamorphism and deformation favor an early Devonian age for the beginning of accretion. The basement of Chilenia underlies the Frontal and Main Cordillera. Extensive upper Paleozoic granitoids on both sides of the Andes (Nasi et al., 1985) and 87Sr/86Sr ratios require an old continental crust (Mpodozis and Kay, 1992). This basement is exposed in small erosional windows and roof pendants in the Cordón del Plata and Cordón del Portillo (Caminos, 1965). Zircons in these rocks have yielded U/Pb ages of 1069 ± 36 Ma (Ramos and Basei, 1997) that are consistent with Laurentian affinities for Chilenia. The final amalgamation of Chilenia occurred in the Early Carboniferous as sedimentary deposits overstepped both terranes. Continental deposits are widespread in the Eastern Precordillera, nearshore marine facies occur in the Western Precordillera, and estuarine to turbiditic facies are in the Frontal and Principal Cordilleras.

Frontal and main Andean Cordilleras Carboniferous to Triassic Subduction, Deformation, and Rifting Subduction-related magmatism resumed in the Late Paleozoic more than 300 km east of the Paleozoic arcs (Ramos et al., 1986) and then ceased in the mid Permian at the time of the strong deformation of the San Rafael orogenic phase (Ramos, 1988a). Both the San Rafael deformation and cessation of subduction-related magmatism have been attributed to the collision of an unidentified terrane (Equis) in the final stages of the amalgamation of Pangea by Mpodozis and Kay (1992). Martínez et al. (2006) argue for a link with shallow subduction to explain these events. After the San Rafael deformation, generalized extension occurred during a period in which the Gondwana plate was stationary with respect to the South Pole (Valencio et al., 1983; Ramos, 1988a). Magmatism in this period produced batholiths like the Colangüil in San Juan Province that was emplaced between 264 and 247 Ma (Llambías and Sato, 1990) and rhyolitic volcanism within the Paleozoic accreted terranes (Kay et al., 1989). This extensive acidic volcanism has been interpreted as evidence of generalized extension during the Triassic (Zeil, 1981) in a period of post-collisional volcanism associated with slab breakoff in the final stages of the amalgamation of Pangea (Mpodozis and Kay, 1992). Alternatively, the rhyolites have been interpreted as being related to Late Permian steepening of a shallow subduction zone (Martínez et al. 2006). ANDEAN HISTORY Most of the early-middle Mesozoic history of the Andes was dominated by an extensional regime closely linked to a negative trench rollback velocity associated with the early stages of the opening of the South Atlantic (Ramos, 1999). The beginning of the drift phase of the Atlantic produced a change to a positive overriding velocity of the South American plate during the middle Cretaceous and the beginning of the contractional history of the Andes. The arc magmatic history includes an eastward shift of the frontal arc from the Cordillera de la Costa in Chile in the Jurassic to the Cordillera Principal in the Miocene and a period of intra-arc extension in the Cretaceous (Mpodozis and Ramos, 1990). The Tertiary history of the transect starts in the Oligocene in a period of localized extension coincident with important volcanism (Godoy et al., 1999; Charrier et al., 2002). This extension ended at ca. 20 Ma when the volcanic activity of the Farellones Formation began (Munizaga and Vicente, 1982). Several interrelated features demonstrate the shallowing of the Benioff zone that followed. The first is an eastward expansion of subductionrelated magmatism from the Main Cordillera in the early Miocene to the eastern Sierras Pampeanas by the latest Miocene to Pliocene (Kay et al., 1987; Ramos et al., 1991, 2002; Kay and Mpodozis, 2002). The geochemical characteristics of the arc magmas in the Main Cordillera indicate that thickening of the continental crust

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started at 18 Ma in relation to tectonic stacking of the Andean Cordillera (Kay et al., 1991). The second feature is the eastward shifting of orogenic deformation during the past 20 Ma with the front at 275 km from the trench at 20–10 Ma, 325 km at 10–5 Ma, and 365 km at 2 Ma. This implies an average propagation rate of 2.5 mm/yr of the orogenic front since 20 Ma, although the shifting was probably episodic (Ramos et al., 2002). The Cenozoic sedimentary history records the eastward migration of the orogenic front. Thick sequences of continental deposits (Santa María Conglomerates of Schiller, 1912) with ages of 20–10 Ma unconformably overlie pre Tertiary rocks in the Cordillera Principal. The angular unconformity is well seen on the field trip route at Cerro Penitentes. Tertiary foreland deposits east of the High Cordillera are distal fluvial facies. A late Miocene unconformity that separates the 10–5 Ma La Pilona beds in the Uspallata valley and Cacheuta region from older Tertiary beds has been associated with the uplift of the Cordillera Frontal (Ramos et al., 1996a, 1996b). Late Miocene and Pliocene deposits in the Uspallata and Cacheuta regions were folded and thrusted before the deposition of the Plio-Pleistocene Mogotes Formation alluvial fan deposits at the present active front in the outer foothills of the city of Mendoza city. Seismic sections and drilling in the plains east of the Precordillera clearly show that the present orogenic front is composed of a set of imbricated thrusts. The thrust front in the eastern side of the Precordillera is still active. Intense contractional deformation can be seen in the field at Sierra de las Peñas (Cortés, 1990; see also chapter 4 of this volume) and inferred from earthquake focal mechanisms and escarpments on alluvial fans. The fanglomerates of the Mogotes Formation and other younger alluvial fans are deformed by this neotectonic activity. FIELD TRIP LOG Drive along the Thrust Front of the Precordillera The field trip begins in the city of Mendoza, which is built along the trace of the Cerro La Cal fault. The fault scarp can be seen across Las Heras Avenue, a few meters west where it intersects Perú Street. The railroad seen here was built along the fault scarp, a few years after the 1861 earthquake that destroyed the city of Mendoza. Leaving Mendoza, Highway 7 to Chile runs parallel to the orogenic front, which is defined at these latitudes by an imbricate fan of thrust sheets with neotectonic activity that deformed and uplifted a complex system of terraces. Stop 1: Aguada Pizarro: Active Thrust Front The first stop is along the International Highway 7 to Chile at km 1073 near 33°04′50.5″S; 69°04′55.8″W. This stop is near the eastern border of the Precordillera where it coincides with the active thrust front of the central Andes at these latitudes. The earthquake, which partially destroyed the southern part of the city of Mendoza in 1985, had

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a hypocenter a few kilometers southwest of the city. The earthquake, which had a magnitude of 5.75 and a compressive focal mechanism, is thought to have occurred along this thrust front. The epicenter is put at 14 km along a reverse fault dipping at 56° to the west (Triep, 1987). The Aguada Pizarro fault, which is part of the active thrust front, can be seen at this stop. The fault is located a few hundred meters to the west of the YPF well PAPx-1, which intersects the base of the Miocene sequence at a depth of 4375 m. Looking to the north, several imbricate east-verging thrusts sequentially thrust the graywackes and shales of the Devonian Villavicencio Formation over the synorogenic deposits of the Miocene Marino and Pliocene Mogotes Formations, which in turn are thrust over early Pleistocene alluvial fan deposits (see Figs. 3 and 4). Some terraces are also cut and displaced by younger Quaternary faults. The general geometry of the faults is controlled by the compressional tectonic inversion of a Triassic rift system. Farther to the north, a Miocene subvolcanic dacitic body can be seen. This is one of the easternmost occurrences of early Miocene retroarc volcanics at these latitudes (e.g., Kay et al., 1991). Subsequently, younger magmatic rocks erupted even farther to the east as the Benioff zone shallowed in the late Miocene and Pliocene.

stage of the Triassic rift. They have an onlap relation with the Permo-Triassic Choiyoi volcanic rocks. Covering the Triassic sediments are the synorogenic deposits of the Miocene Mariño Formation (14–9 Ma), which is associated with the main uplift of the Main Cordillera at these latitudes (Fig. 5). The Cuyo basin is located to the east of an extensive sequence of Ordovician oceanic mafic magmatic rocks that mark the boundary between the Cuyania terrane which includes the Precordillera, and the Chilenia terrane to the west. These terranes were most likely sutured together in the Late Devonian to the Early Carboniferous (Ramos et al., 1986). The western margin of the Cuyo basin parallels the proposed terrane suture for more than 700 km along the boundary of the upper plate. The Ischigualasto and the Beazley basins farther east have a similar setting in that they developed on the eastern side of the Cuyania terranes near its Late Ordovician to Early Silurian suture with the Pampia terrane, which incorporates the Sierras Pampeanas east of the Pie de Palo range (see Ramos, 2004). The road continues west and crosses the Río Blanco valley, which coincides with an inverted normal fault that puts Triassic sediments over Miocene deposits. Farther west, the thick sequence of black shales is again the continental lacustrine deposits of the Triassic Cacheuta Formation, which is the main source rock for petroleum in the prolific Cuyo Basin.

Drive across the Southern End of the Precordillera The road continues to the west and north. Along the northern side is the Sierra de Cacheuta, which limits the southernmost part the Precordillera range. The general south-plunging structure in this region is related to the southern boundary of the flat-slab segment. To the south, the Precordillera disappears and is replaced by the mildly structurally inverted Triassic Cuyo rift basin. The Sierra de Cacheuta is composed of Devonian sedimentary rocks that are intruded by Carboniferous to Early Permian diorites and granites. This sequence is unconformably overlain by Permo-Triassic Choiyoi volcanic rocks and Triassic rift deposits. Sedimentary sequences related to the passive eastern side of the Triassic rift can be seen on both sides of the road. These rocks are principally black shales that formed in the sag

Stop 2: Triassic Cuyo Rift Basin and Miocene Foreland Basin Stop along the international road 7 to Chile at 32°56′09.9″; 69°12′58.5″; 1400 m above sea level (asl). The overlap of Triassic rift basins in southern South America with the Permo-Triassic Choiyoi granite–rhyolite province suggests a genetic relation in an extensional regime. The Choiyoi rocks are crustal melts that formed in association with extensive basaltic underplating during a period of relatively slow motion of the Gondwana supercontinent over the underlying mantle. The Cuyo rift basin developed slightly east of the main Choiyoi magmatic belt and is partially coeval with the termination of the acidic magmatism. The margins of the Triassic basins follow

Figure 3. Thrust front at Aguada Pizarro. Fanglomerates of Late Pliocene Mogotes Formation are in fault contact with Early Pleistocene deposits.

Potrerillos Dam

Cenozoic deposits

Paleozoic rocks

Río Men doza

Paleozoic rocks

33°S-

Aguada Pizarro section Figure 4B

to Potrerillos

to Mendoza city

STOP 1

TRIASSIC CUYO BASIN

5 km

A

69°W

AGUADA PIZARRO STRUCTURAL SECTION

Aguada Pizarro well STOP 1 (Figure 3)

Villavicencio Formation

B

Mariño Formation

4375 m

Mogotes Formation

Late Pleistocene deposits

1 km

Villavicencio Formation

Figure 4. (A) Geologic map of the orogenic front of the Precordillera based on Rebori (1979). (B) Cross section of Stop 1 along the tectonic contact of a secondary Pleistocene fault.

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Figure 5. Geologic map of the Potrerillos region west of Mendoza (based on Caminos, 1965).

first order tectonic boundaries (Fig. 5) with the hanging-walls following the suture zones of terranes that were accreted during the Paleozoic. Opposing polarities of half-grabens developed a local transtensional regime with transfer zones linking rift segments. Brittle faulting that took place during the rift stage allowed basaltic magmas to penetrate the cooling, refractory crust (Ramos and Kay, 1991). Olivine-bearing alkali basalts of Middle Triassic (235 Ma) age are interfingered with late synrift deposits. Relatively low degrees of melting (4%–5%) in the mantle are suggested for the basalts, consistent with the comparatively narrow width of the Cuyo basin and their eruption during the last stages of Choiyoi magmatism (Ramos and Kay, 1991). No evidence of volcanic activity is found in the Cuyo rift during most of the Late Triassic, as this period was dominated by a generalized subsidence related to the thermal decay and sedimentary loading that characterizes sag phases. The upper crust became more brittle as silicic magmatism terminated allowing brittle fracturing of the upper crust to create a pathway for relatively primitive basaltic magmas to reach the surface. The basalts could represent a late phase in the stationary supercontinent-related heating of the upper mantle. Subsequently, the Middle Jurassic Andean arc developed to the west contemporaneous with generalized backarc rifting associated with the early opening of the South Atlantic. Arc volcanism is mainly concentrated on the western slope of the Main Andes. The synrift deposits of the Triassic Río Mendoza Formation fanglomerates are seen at Stop 2 (Fig. 6). They are separated by a strong angular unconformity from the Early Paleozoic

rocks of the Villavicencio Formation, and the volcanics of the Choiyoi Group. The conglomerates of the Rio Mendoza Formation represent the alluvial fan deposits that form the basal synrift sequence of the Cuyo rift. Above them, the rest of the sequence of the Triassic rift deposits in the Uspallata Group can be seen. Immediately above the Río Mendoza Formation are the red sandstones and shales and interbedded fine conglomerates that compose the characteristic fluvial facies of the overlying Potrerillos Formation. These two formations constitute the first cycle of the Triassic rift sequence. Above them are the lacustrine deposits of the Cacheuta Formation that mark the beginning of the sag phase. The black shales of the Cacheuta Formation are the most important source rock of the oil fields of the Cuyo basin, southeast of the city of Mendoza. Synorogenic deposits of mainly Miocene age unconformably overlie the Triassic rift deposits. These beds represent the foreland basin distal facies of the proximal Santa María Conglomerates, which will be seen at the stop at Cerro Penitentes. Drive across the Cordillera Frontal Route 7 continues west into the Cordillera Frontal, which is the locus of the Late Paleozoic tectonic activity that has been called the Gondwanian orogeny since the early work of Keidel and Du Toit. The Frontal Cordillera contains the early Permian subduction-related magmatic rocks as well as the subsequent widespread Permo-Triassic rhyolitic magmatic rocks known as the Choiyoi province (Kay et al., 1989).

Frontal and main Andean Cordilleras

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Figure 6. Synrift deposits of the Río Mendoza and Potrerillos Formations at Stop 2.

These Late Paleozoic–Triassic batholiths and associated acidic volcanic rocks record the history of an important period of crustal growth on the margin of western Gondwana. These events follow a period of Middle to Late Paleozoic subduction that culminated with uplift and crustal thickening, followed by extensional collapse of the orogen associated with extensive melting of the crust (Mpodozis and Kay, 1992). The formation of the Choiyoi province has been associated with a relatively stable period in the history of Gondwana during which underplated mantle basalts caused extensive crustal melting. There is no evidence for subduction of oceanic crust beneath the Cordillera Frontal during this time. There are synextensional faults in the Choiyoi group sequences. These events are related to the final assembly of the Gondwana continent and the first stages of its fragmentation. Late Paleozoic subduction of the proto-Pacific Ocean along the new western border of Gondwana resulted in formation of the wide accretionary prism that is now exposed along the Chilean coast and the plutonic and volcanic rocks known as the pre-Choiyoi assemblage (Kay et al., 1989; Mpodozis and Kay, 1990) in the Cordillera Frontal. Synchronous intra-arc and retroarc sedimentary basins forming in both Argentina and Chile were filled by several thousand meters of Devonian (?) to Lower Permian marine and continental sediments (Caminos, 1979; Polanski, 1970). The Late Permian to Triassic history of the Cordillera Frontal region is dominated by the extensive Choiyoi granite-rhyolite province whose formation precedes and in part overlaps the development of Triassic rift basins just to the east (Ramos and Kay, 1991). The Choiyoi province extends for more than 2500 km from Collaguasi (22°S) in northern Chile to the Neuquén Basin and northern Patagonian Andes near 40°S (Mpodozis and Kay, 1990). Particularly extensive sequences of rhyolitic ignimbrites (Cortés, 1985) and granitic batholiths occur (e.g., Colangüil and San Guillermo batholiths; Llambías and Sato, 1990) occur in the San Rafael block in Argentina to the south (34°S) and in the Cordillera Frontal (29°–33°S; Caminos, 1979). Volcanic calderas are still recognizable in some places. A major lower Permian compressional deformation is recognized in the Cordillera Frontal where a pronounced uncon-

formity separates folded and faulted Devonian and Carboniferous sediments from undeformed Middle to Upper Permian Choiyoi volcanic sequence (Fig. 7). This deformational phase known as the San Rafael (Polanski, 1970) was also recognized as a period of low magmatic activity in western Argentina (Pérez and Ramos, 1990). Farther east, in the Carboniferous Paganzo basin, in the modern Precordillera and Sierras Pampeanas, no deformation is observed. At that time, a change occurs from deposition of active rifting of synrift facies to sag facies. One possible explanation for compressional deformation and crustal thickening is that oblique collision of an exotic terrane occurred along the coast at this time. This scenario is consistent with the work of Rapalini (1989), who used paleomagnetic evidence to suggest that pre–mid-Permian rocks in Argentina are rotated as a consequence of oblique plate motion along the Pacific margin. Late Permian rocks are unrotated. The orientation and intensity of the mylonitic and cataclastic zones in the granitoids are also most easily explained by oblique motion (Mpodozis and Kay, 1990). Collision could also explain the apparent end of subduction related magmatism. An alternative explanation was presented by Martínez et al. (2006) who related the San Rafael deformation and the cessation of arc related magmatism to a period of flat-subduction. Between Stops 3 and 4, post-tectonic granitoids such as the Guido Granite (Fig. 7) can be seen in the western part of Precordillera. These granites were emplaced during the Late Paleozoic. They have an Early Permian age and could be subduction-related granitoids similar to the ones described in Chile farther north (Mpodozis and Kay, 1990). Mafic dikes were emplaced during the late episodes of crystallization. The post-collisional granites and rhyolites typical of the Choiyoi province show chemical characteristics consistent with an origin by crustal melting with heat provided by basaltic underplating. The Choiyoi province is only one of the granite-rhyolite provinces that formed along the western border of Gondwana from the Upper Paleozoic to the Lower Jurassic (see Mpodozis and Kay, 1990). Others include the Mitu province between Peru and Bolivia, the Choiyoi province discussed here, the Chon Aike province in Patagonia, the Antarctica Peninsula and Ellsworth-

Figure 7. Drive between Stops 2 and 4 through the Cordillera Frontal where the main types of the Late Paleozoic magmatic rocks can be observed (based on Caminos, 1965).

Frontal and main Andean Cordilleras Witmore in Antarctica, and the New England fold belt in Australia. The combined outcrops form a belt of more than 10,000 km, developed along the Pacific margin of Gondwana, preceding its rupture and dispersal. In most of these regions, Late Paleozoic calc-alkaline, arc-related plutons are deformed and uplifted. These events are followed by the intrusion of high-level hypersilicic rhyolites and granites ranging in age from Upper Permian to Lower Jurassic. In all cases, these events represent the final amalgamation of Gondwana, the extensional collapse of compressional regimes and the initial stages of breakup of the supercontinent. The tectonic processes occurring are those associated with subduction, terrane collision, and extension. The large amounts of granite and rhyolite generated are the results of the focusing of these events during a short period of time, associated with the formation and breakup of a supercontinent. Stop 3: Las Carreras Fault and the Permo-Triassic Choiyoi Volcanics Stop is along Highway 7 at 32°52′0.3″; 69°16′10″; 1523 m asl. The Las Carreras fault seen at this stop is the northern continuation of the Cordón del Plata thrust front. In this location, the Permian Guido Granite is thrust over Carboniferous sedimentary deposits and the Choiyoi volcanic rocks (Figs.7 and 8). The fault was responsible for the late Miocene uplift of the Cordillera Frontal and was active during late Miocene and early Pliocene times, as inferred from the synorogenic Tertiary deposits east of the Cordón del Plata just to the south (Ramos, 1993). The volcanic rocks of the Choiyoi Group seen in this region can be separated into two sections. The lower section, which can be seen at km 1107 along Highway 7, is formed by andesitic volcanic rocks that represent the pre-Choiyoi stage of subduction related magmatism. The upper sequence is dominated by siliceous rhyolites. Stop 4: Southern End of the Uspallata Valley Stop is along Highway 7 at 32°36′53.5″; 69°23′54.1″; 1874 m asl.

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The southern end of the Calingasta-Uspallata valley is seen at this stop (Figs. 7 and 9). At these latitudes, the Cordillera Frontal is thrust over the Precordillera. A series of imbricate thrust put Choiyoi volcanic rocks repeatedly on top of Miocene continental foreland basin strata. Drive across the Cordillera Principal: The High Andes The Cordillera Principal or High Andes starts west of Punta de Vacas. Here the Gondwana evolutionary stage of the Cordillera is heavily overprinted by Mesozoic and Cenozoic events. The Mesozoic to Cenozoic evolution is characterized by two stages: (1) the Jurassic to Early Cretaceous intra-arc development of a carbonatic sedimentary basin that is intermittently interrupted by volcanic activity (Fig. 10); and (2) the formation of a Cenozoic fold-and-thrust belt. Tertiary deformation and contraction started in Chile in the Paleogene at these latitudes with an extensional stage during the Oligocene (Godoy et al., 1999). The structures seen in the Aconcagua fold-and-thrust belt on the Argentine side of the Cordillera are mainly early Miocene to late Miocene in age (Fig. 11). One of the outstanding features of the first stage is the development of Mesozoic marine sequences that were controlled by transgressions and regressions from the Pacific side. Those sequences are grouped into four sedimentary cycles, which are separated by regional first-order unconformities. These sequences are clearly depicted in the Neuquén Basin farther to the south, where a well-developed Liassic to Neocomian marine retroarc basin developed behind the magmatic arc in the eastern foothills of the Cordillera. The Neuquén Basin is linked to the north with the Aconcagua basin, which we are seeing here. The Aconcagua Basin has a different paleogeography with a larger participation of volcanic rocks, but similar middle Jurassic to Cretaceous stratigraphic cycles. There are interesting parallels between the evolution of the basin and the magmatic arc history (Ramos, 1985a, 1985b). The main periods of regional unconformities in the foreland are coincident with periods when the magmatic arc migrated eastward. The intermittent nature of the magmatic activity, as well as the spatial variation of the volcanic front in the Andes, are closely related to changes in plate motion

Ca

rb

on

ife

ro

us

de

po

sit

s

Figure 8. View to the south of Las Carreras fault. Note the imbrication of a thin slice of Carboniferous deposits between the rhyolites of the Choiyoi Group and the Permian granite.

Ramos

controlled by variations in the spreading velocities of the Pacific and Atlantic oceanic ridges. Removal of the continental margin by forearc subduction erosion can also be related to the migration of the arc volcanic fronts. Such process can be seen to be combined with the shallowing of the subduction zone at this latitude and just to the south in the past 10 Ma (see Kay et al., 2005). Most of the Chilean and the westernmost Argentine Mesozoic basins of the Cordillera Principal are intra-arc basins controlled by the development of two distinctive arcs: an inner more active arc along the Cordillera de la Costa of Chile concentrates the main andesitic activity, and an outer arc produces rock suites that are mainly of andesitic to bimodal composition. Several authors have proposed an extensional regime in the arc massif region mainly during the Early Cretaceous, which was responsible for the intra-arc basin development. This process has been envisaged as an intracontinental spreading which has controlled the rapid subsidence of the volcanic pile where burial metamorphism closely followed extrusion (Levi and Aguirre, 1981). This extensional regime in the arc has been attributed to a negative trench rollback velocity prior to the opening of the South Atlantic Ocean of the Western Gondwana plate. The opening of the South Atlantic Ocean in the Neocomian (ca. 125 Ma) changed the absolute motion of South America to a positive trench rollback velocity and to a compressive regime in the arc (Ramos, 1999). The intermittent activity of the outer arc is closely related to periods of sea-level lowstands in both the intra-arc and the retroarc basins. These local sea-level changes have been partially correlated with the worldwide eustatic onlap cycles. The intra-arc basins were active until Early Barremian times when an important eastward migration of the main magmatic arc occurred, together with a low stand period in the retroarc and intra-arc basins and the development of a single and expanded central arc. The retroarc easternmost basin was exclusively continental from this time on, and the Pacific sea no longer reached the eastern side of the cordillera. This important paleogeographic change is closely linked with the beginning active spreading in the South Atlantic. Soon after the opening of the South Atlantic, a new increase in plate motion started a compressive regime, which eliminated the intra-arc basins and the arc front moved toward the foreland. This second step is recognized in these latitudes at ca. 110 Ma when the Cristo Redentor and the Juncal Formations were deposited. The eastward expansion of the magmatic arc could be related to a decrease in the subducted slab dip, coeval with a higher convergence rate, combined by the tectonic erosion at the subduction zone. The maximum sea-floor spreading rate at the Pacific and the South Atlantic spreading centers was reached in the Late Cretaceous (ca. 80 Ma). A positive trench rollback velocity at this time may be responsible for the development of a fold-and-thrust belt on the eastern flank of the Andean orogen. At the same time, a foreland basin formed at the leading edge of the deformational front in response to tectonic loading in the adjacent thrust belt. At the final stage of compression, several granitoid stocks were emplaced in the arc massif.

Figure 9. Imbrication of Choiyoi volcanics with Miocene synorogenic deposits at the southern end of the Uspallata valley. View to the south. Note the remnants of the Miocene peneplain on the western side of the road.

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Frontal and main Andean Cordilleras

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Figure 10. Intra-arc extension developed in the Cordillera Principal during the Late Jurassic–Early Cretaceous (after Ramos, 1989).

Figure 11. The Aconcagua fold-and-thrust belt (FTB) at Puente del Inca (after Ramos, 1985b).

Stop 5: Aconcagua Fold-and-Thrust Front: Cerro Penitentes Stop is along Highway 7 at 32°49′56.3″; 69°52′17.7″; 2654 m asl. Stop 5 shows the thrust front of the Miocene Aconcagua fold-and-thrust belt (Fig. 11). Middle to Late Jurassic limestones (La Manga Formation) are seen to be thrust over thick early to

middle Miocene Santa María conglomerates that in turn unconformably overlie marine and continental Jurassic strata. The eastverging Penitentes thrust dips at 5° to 22° to the west. The effects of erosion have led to the near formation of a klippe on the top of Cerro Visera farther east. A thin blanket of Choiyoi volcanic rocks unconformably overlies Carboniferous deposits and Permian granitoids, where Late Permian–Early Triassic extensional faults are preserved (Fig. 12).

Figure 12. The thrust front of the Aconcagua fold-and-thrust belt at Penitentes.

Figure 13. Panoramic view of Puente del Inca section with the present structural interpretation. Crt—Alto Tupungato Formation (Carboniferous), Trch—Choiyoi Group (Triassic), Jm—La Manga Formation (Middle Jurassic), Jt—Tordillo Formation (Late Jurassic), JKi—Early TithoNeocomian deposits, JKm—Middle Neocomian deposits, JKs—Late Neocomian deposits, Tc—Santa María Conglomerates (Miocene), Tv— Puente del Inca trachyte (15 Ma, Miocene).

Frontal and main Andean Cordilleras Stop 6: Puente del Inca: Aconcagua Fold-and-Thrust Belt Stop 6 is along Highway 7 at pullout for the Puente del Inca. There is a hostería here as well as tourist shops and places to buy food. The view to the south from this stop is the most classic section of the High Andes. Compare the evolution of the knowledge since Darwin and Schiller (see figures 2 and 3 in chapter 4 of this volume) with the present interpretation of Figures 13 and 14. The section shows the autochthon, which is composed of Carboniferous hornfels, thin pyroclastic deposits of Choiyoi Group and Jurassic limestones and conglomerates of proximal facies, and small outcrops of marine Early Cretaceous; the first thrust plate, which is composed of Middle to Late Jurassic marine and

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continental deposits and Early Cretaceous continental to transitional marine deposits; and the second thrust plate, which has Late Jurassic gypsum at the base and Titho-Neocomian continental and marine strata above. All these sequences are intruded by thick Miocene trachytic dikes (15 Ma). Cretaceous basaltic and andesitic lenses that are interbedded in the sequence in the first thrust sheet become more abundant in the westernmost thrust sheets (Ramos, 1985a, 1985b). The Puente del Inca is a natural bridge formed during the last deglaciation, probably in the Late Pleistocene, as an ice bridge associated with an avalanche. Biogenic sulfates and carbonates precipitated from thermal springs located along the Penitentes thrust cemented the avalanche deposits forming the present bridge (Ramos, 1993).

Figure 14. Geologic map of the Aconcagua fold-and-thrust belt. Cat—Carboniferous Alto Tupungato Formation; Trch—Volcanics of the Choiyoi Group; c—Early Jurassic breccia; Jlm—Jurassic La Manga Formation; e—Gypsum; of Auquilco Formation; Jt—Red beds of Jurassic Tordillo Formation; JKvm—Tithono-Berriasian Vaca Muerta Formation; Km—Sandstones of Valanginian Mulichinco Formation; Ka—Limestones of Hauterivian Agrio Formation; Kd—Red beds of Barremian Diamante Formation; Kv—Early Cretaceous volcanics; l—Miocene dacites of Puente del Inca; m—Early Cretaceous diorites; Tsm—Miocene Santa María Conglomerates; Qgl and Qsl—glacial and slump deposits; Q— undifferentiated alluvial sediments.

Figure 15. View to the south of Mount Aconcagua in Valle de los Horcones. Thrust front 2, which coincides with this valley, covered by glacial moraines and avalanches.

Figure 16. View to the south at Stop 7, where thrust faults two (Quebrada Blanca), three (La Yesera), and four (Las Leñas) repeat different facies of the Late Jurassic to Early Cretaceous deposits. These faults are attached in the Auquilco Gypsum.

Figure 17. Location of the main thrust sheets of the Aconcagua fold-and-thrust belt.

Frontal and main Andean Cordilleras Stop 7: Highest Mountain of the Americas: Cerro Aconcagua (6967 m asl) Stop is along Highway 7 at Mirador (view point) Aconcagua near 32°49′22″; 69°56′17.5″; 2847 m asl. Climb above the highway for a good view. At this stop there is a magnificent view of the Pared Sur (south wall) of the Aconcagua (6967 m asl). The wall is formed by volcanic and breccia flows of andesitic composition of the Aconcagua Volcanic Complex (15–8.9 Ma, Ramos and Yrigoyen, 1987; Godoy et al., 1988; Kay et al.., 1991). On the western side of the valley (Fig. 15), there is an imbrication of Jurassic continental red beds and Early Cretaceous limestones. On the eastern side, the diapiric effects of the Late Jurassic gypsum of the Auquilco Formation produced the complex structure of Cerro Panta. To the south (Fig. 16), the fault imbrications are visible from the second to the fourth thrust sheets repeating the different Titho-Neocomian units.

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normal thickness of these units, indicating their proximity to the volcanic arc. A large rock avalanche was produced from one of the volcaniclastic members of the Mesozoic deposits. Stop 8a: Optional Stop (Darwin Shelter) Depending on the time available and the weather, an optional stop can be made to visit a colonial shelter built in 1765 (Fig. 19). The Spanish Royal Mail constructed those shelters to cross the Main Andes during the winter, because at that time due to the war with the British, they did not have safe access to cross the Strait of Magellan.

A

Stop 8: Cretaceous Volcanic Arc near Las Cuevas along the Chilean Border Stop is at 32°40′49.0″; 70°03′14.9″; 3224 m asl. Stop 8 shows that the westernmost thrust sheets are dominated by volcanic and volcaniclastic rocks derived from the volcanic arc (Figs. 17 and 18). These Early Cretaceous (Neocomian) rocks have Valanginian (137–132 Ma) ammonites, which corroborate the Neocomian age of the sequence (see Aguirre-Urreta and Rawson, 1997). The thrusts formed an imbricate fan with the western and older thrusts rotated 25° to vertical and almost overturned. In the northern margin, the Neocomian limestones override Late Jurassic–Early Cretaceous red beds. The thickness of these continental and volcaniclastic deposits exceeds several times the

Figure 18. The limestones and shales of the Cretaceous Agrio Formation bear Olcostephanus atherstoni, a key fossil of widespread distribution.

B

Figure 19. Present view of (A) Las Vacas and (B) Paramillos de las Cuevas shelters.

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We are going to visit the Paramillo de las Cuevas shelter (“casucha”), which has a prolific history and was built by the Spaniards as one of eight shelters across the Andes. Several famous travelers described the penuries related to the winter and the snow storms and strong winds (see Darwin stories in his famous trip across the Andes in 1835). ACKNOWLEDGMENTS This field guide has been modified from that originally written for the one-day intra-conference field trip on Wednesday, 5 April 2006, during the Backbone of the Americas meeting in Mendoza, Argentina. Thanks to the seven other leaders and co-leaders who made the trip possible for 380 participants on the original trip. Leaders and co-leaders included Suzanne Mahlburg Kay (Cornell University, USA), Constantino Mpodozis and Pamela Alvarez (SiPetrol, Chile), Estanislao Godoy (Servicio Nacional de Geología y Minería [SERNAGEOMIN], Chile), Laura Giambiagi and Florencia Bechis (Centro Regional de Investigaciones Científicas y Técnicas [CRICYT], Mendoza, Argentina), Sergio Orts (Rio Tinto, Argentina), and Daniel Pérez, Silvia Barredo, Andrés Folguera, Daniel Yagupsky, Maximiliano Neipauer, Víctor García, and Fernando Pose (Laboratorio de Tectónica Andina, Universidad of Buenos Aires, Argentina). REFERENCES CITED Aguirre-Urreta, M.B., and Rawson, P.F., 1997, The ammonite sequence in the Agrio Formation (Lower Cretaceous), Neuquén basin, Argentina: Geological Magazine, v. 134, p. 449–458, doi: 10.1017/ S0016756897007206. Astini, R.A., Benedetto, J.L., and Vaccari, N.E., 1995, The early Paleozoic evolution of the Argentina Precordillera as a Laurentian rifted, drifted, and collided terrane: A geodynamic model: Geological Society of America Bulletin, v. 107, p. 253–273, doi: 10.1130/0016-7606(1995)107 <0253:TEPEOT>2.3.CO;2. Astini, R.A., Ramos, V.A., Benedetto, J.L., Vaccari, N.E., and Cañas, F.L., 1996, La Precordillera: Un terreno exótico a Gondwana: XIII Congreso Geológico Argentino y III Congreso Exploración de Hidrocarburos, Actas 5, p. 293–324. Benedetto, J.L., and Astini, R., 1993, A collisional model for the stratigraphic evolution of the Argentine Precordillera during the Early Paleozoic: Paris, Second Symposium International Géodynamique Andine, ISAG 93 (Oxford), p. 501–504. Benedetto, J.L., Sánchez, T.M., Carrera, M.G., Brussa, E.D., and Salas, M.J., 1999, Paleontological constraints in successive paleogeographic positions of Precordillera terrane during the Early Paleozoic, in Ramos, V.A., and Keppie, D., eds., Laurentia-Gondwana connections before Pangea: Geological Society of America Special Paper 336, p. 21–42. Bond, G.C., Nickelson, P.A., and Kominz, M.A., 1984, Breakup of a supercontinent between 625 Ma and 55 Ma: New evidence and implications for continental histories: Earth and Planetary Science Letters, v. 70, p. 325–345, doi: 10.1016/0012-821X(84)90017-7. Borrello, A.V., 1969, Los geosinclinales de la Argentina: Dirección Nacional de Geología y Minería Anales 14, 136 p. Cahill, T., and Isacks, B.L., 1992, Seismicity and the shape of the subducted Nazca plate: Journal of Geophysical Research, v. 97, p. 17,503–17,529. Caminos, R., 1965, Geología de la vertiente oriental del Cordón del Plata, Cordillera Frontal de Mendoza: Asociación Geológica Argentina Revista, v. 20, p. 351–392. Caminos, R., 1979, Cordillera Frontal, in Turner, J.C.M., ed., Segundo simposio de geología regional Argentina: Córdoba, Academia Nacional de Ciencias 1, p. 397–454.

Caminos, R., ed., 2000, Geología Argentina: Buenos Aires, Instituto de Geología y Recursos Minerales, Anales 29, p. 1–796. Charrier, R., Baeza, O., Elgueta, S., Flynn, J.J., Gans, P., Kay, S.M., Munoz, N., Wyss, A.R., and Zurita, E., 2002, Evidence for Cenozoic extensional basin development and tectonic inversion, south of the flat-slab segment, southern central Andes, Chile (33°–36°S.L.): Journal of South American Earth Sciences, v. 15, p. 117–140, doi: 10.1016/S0895-9811(02)00009-3. Chinn, D.S., and Isacks, B., 1983, Accurate source depths and focal mechanisms of shallow earthquakes in western South America and in the New Hebrides island arc: Tectonics, v. 2, p. 529–564. Cortés, J.M., 1985, Vulcanitas y sedimentitas lacustres en la base del Grupo Choiyoi al sur de la estancia Tambillos, provincia de Mendoza, República Argentina: Antofagasta, IV Congreso Geológico Chileno, Actas 1, p. 89–108. Cortés, J.M., 1990, Estudio geológico estructural de la Sierra de las Peñas: Servicio Geológico Nacional (unpublished). Dalla Salda, L., Cingolani, C., and Varela, R., 1992a, Early Paleozoic orogenic belt of the Andes in southwestern South America: Result of Laurentia-Gondwana collision?: Geology, v. 20, p. 617–620, doi: 10.1130/ 0091-7613(1992)020<0617:EPOBOT>2.3.CO;2. Dalla Salda, L., Dalziel, I.W.D., Cingolani, C.A., and Varela, R., 1992b, Did the Taconic Appalachians continue into southern South America?: Geology, v. 20, p. 1059–1062, doi: 10.1130/0091-7613(1992)020<1059:DTTACI> 2.3.CO;2. Dalziel, I.W.D., 1991, Pacific margins of Laurentia and East Antarctica-Australia as a conjugate rift pair: Evidence and implications for an Eocambrian supercontinent: Geology, v. 19, p. 598–601, doi: 10.1130/0091-7613 (1991)019<0598:PMOLAE>2.3.CO;2. Dalziel, I.W.D., 1997, Neoproterozoic-Paleozoic geography and tectonics: Review, hypothesis, environmental speculation: Geological Society of America Bulletin, v. 109, p. 16–42, doi: 10.1130/0016-7606(1997)109 <0016:ONPGAT>2.3.CO;2. Dalziel, I.W.D., Dalla Salda, L., Cingolani, C., and Palmer, P., 1996, The Argentine Precordillera: A Laurentian terrane?: GSA Today, v. 6, p. 16–18. Darwin, C., 1846, Geological observations of South America, being the third part of the Geology of the voyage of the Beagle during 1832–1836: London, Smith, Elder, 279 p. Godoy, E., Harrington, R., Fierstein, J., and Drake, R., 1988, El Aconcagua, parte de un volcán mioceno?: Revista Geológica de Chile, v. 15, p. 167–172. Godoy, E., Yañez, G., and Vera, E., 1999, Inversion of an Oligocene volcanotectonic basin and uplifting of its superimposed Miocene magmatic arc in the Chilean Central Andes: First seismic and gravity evidences: Tectonophysics, v. 306, p. 217–336, doi: 10.1016/S0040-1951(99)00046-3. González Bonorino, F., 1950, Geologic cross section of the Cordillera de los Andes at about parallel 33°S L. (Argentina-Chile): Geological Society of America Bulletin, v. 61, p. 17–26, doi: 10.1130/0016-7606(1950)61 [17:GCOTCD]2.0.CO;2. Handschumacher, D.W., 1976, Post Eocene plate tectonics of the Eastern Pacific, in Sutton, G.H., Manghnani, M.H., and Moberly, R., eds., The geophysics of the Pacific Ocean and its margins: American Geophysical Union, p. 117–202. Isacks, B., 1988, Uplift of the Central Andean plateau and bending of the Bolivian orocline: Journal of Geophysical Research, v. 93, p. 3211–3231. Jordan, T., Isacks, B., Ramos, V.A., and Allmendinger, R.A., 1983a, Mountain building in the central Andes: Episodes, v. 6, p. 20–26. Jordan, T., Isacks, B., Allmendinger, R., Brewer, J., Ramos, V.A., and Ando, C., 1983b, Andean tectonics related to geometry of subducted plates: Geological Society of America Bulletin, v. 94, p. 341–361, doi: 10.1130/ 0016-7606(1983)94<341:ATRTGO>2.0.CO;2. Kay, S.M., and Mpodozis, C., 2002, Magmatism as a probe to the Neogene shallowing of the Nazca plate beneath the modern Chilean flat-slab: Journal of South American Earth Sciences, v. 15, p. 39–58, doi: 10.1016/ S0895-9811(02)00005-6. Kay, S.M., Maksaev, V., Moscoso, R., Mpodozis, C., and Nasi, C., 1987, Probing the evolving Andean lithosphere: Mid-late Tertiary magmatism in Chile (29°–30°30′S) over the modern zone of subhorizontal subduction: Journal of Geophysical Research, v. 92, p. 6173–6189. Kay, S.M., Ramos, V.A., Mpodozis, C., and Sruoga, P., 1989, Late Paleozoic to Jurassic silicic magmatism at the Gondwanaland margin: Analogy to the Middle Proterozoic in North America?: Geology, v. 17, p. 324–328, doi: 10.1130/0091-7613(1989)017<0324:LPTJSM>2.3.CO;2.

Frontal and main Andean Cordilleras Kay, S.M., Mpodozis, C., Ramos, V.A., and Munizaga, F., 1991, Magma source variations for mid to late Tertiary volcanic rocks erupted over a shallowing subduction zone and through a thickening crust in the Main Andean Cordillera (28–33°S), in Harmon, R.S, and Rapela, C., eds., Andean magmatism and its tectonic setting: Geological Society of America Special Paper 265, p. 113–137. Kay, S.M., Godoy, E., and Kurtz, A., 2005, Episodic arc migration, crustal thickening, subduction erosion, and magmatism in the south-central Andes: Geological Society of America Bulletin, v. 117, p. 67–88, doi: 10.1130/B25431.1. Levi, B., and Aguirre, L., 1981, Ensialic spreading subsidence in the Mesozoic and Paleogene Andes of Central Chile: Journal of Geological Society of London, v. 138, p. 75–81. Levi, B., Aguirre, L., and Nystrom, J.O., 1982, Metamorphic gradients in burial metamorphosed vesicular lavas: Comparison of basalt and spilite in Cretaceous basic flows: Contributions to Mineralogy and Petrology, v. 80, p. 49–58, doi: 10.1007/BF00376734. Llambías, E., and Sato, A.M., 1990, El batolito de Colangüil (29–31°S), Cordillera Frontal de Argentina: Estructura y marco tectónico: Revista Geológica de Chile, v. 17, p. 89–108. Martínez, A., Rodríguez Blanco, L., and Ramos, V.A., 2006, Permotriassic magmatism of the Choiyoi Group in the Cordillera Frontal of Mendoza, Argentina: Geological variations associated with changes in paleo-Benioff zone: Mendoza, Backbone of the Americas Conference, Abstracts with Program, p. 60. Moores, E.M., 1991, Southwest U.S.–East Antarctic (SWEAT) connection: A hypothesis: Geology, v. 19, p. 425–428, doi: 10.1130/0091-7613(1991) 019<0425:SUSEAS>2.3.CO;2. Mpodozis, C., and Kay, S.M., 1990, Late Paleozoic to Triassic evolution of the Gondwana margin: Evidence from Chilean Frontal Cordilleran Batholiths (28°–33°S): Revista Geológica de Chile, v. 17, p. 153–180. Mpodozis, C., and Kay, S.M., 1992, Late Paleozoic to Triassic evolution of the Gondwana margin: Evidence from Chilean Frontal Cordilleran Batholiths (28°–31°S): Geological Society of America Bulletin, v. 104, p. 999–1014, doi: 10.1130/0016-7606(1992)104<0999:LPTTEO>2.3.CO;2. Mpodozis, C., and Ramos, V.A., 1990, The Andes of Chile and Argentina, in Ericksen, G.E., Cañas Pinochet, M.T., and Reinemud, J.A., eds., Geology of the Andes and its relation to hydrocarbon and mineral resources: Houston, Circumpacific Council for Energy and Mineral Resources, Earth Sciences Series 11, p. 59–90. Munizaga, F., and Vicente, J.C., 1982, Acerca de la zonación plutónica y del vulcanismo miocénico en los Andes de Aconcagua (lat. 32°–33°S): Datos radiométricos K/Ar: Revista Geológica de Chile, v. 16, p. 3–21. Nasi, C., Mpodozis, C., Cornejo, P., Moscoso, R., and Maksaev, V., 1985, El batolito de Elqui-Limarí (Paleozoico superio-Triásico): Características petrográficas, geoquímicas y significado tectónico: Revista Geológica de Chile, v. 25–26, p. 77–111. Pardo, M., Monfret, T., and Comte, D., 2000, Sismotectónica y distribución de esfuerzos en la zona de subducción subhorizontal de Chile Central (30°–32°S) utilizando datos telesísmicos: IX Congreso Geológico Chileno, Actas 2, p. 459–463. Pardo, M., Comte, D., and Monfret, T., 2002, Seismotectonic and stress distribution in the central Chile subduction zone: Journal of South American Earth Sciences, v. 15, p. 11–22, doi: 10.1016/S0895-9811(02)00003-2. Pardo Casas, F., and Molnar, P., 1987, Relative motion of the Nazca (Farallon) and South American plates since Late Cretaceous time: Tectonics, v. 6, p. 233–248. Pérez, D., and Ramos, V.A., 1990, La actividad magmática gondwánica, Late Paleozoic of South America: Buenos Aires, International Geological Correlation Programme (IGCP) Project 211, Annual Meeting of the working group, Abstracts, p. 89–92. Pilger, R.H., 1981, Plate reconstructions, aseismic ridges, and low angle subduction beneath the Andes: Geological Society of America Bulletin, v. 92, p. 448–456, doi: 10.1130/0016-7606(1981)92<448:PRARAL>2.0.CO;2. Pilger, R.H., 1984, Cenozoic plate kinematics, subduction and magmatism: South American Andes: Journal of the Geological Society, v. 141, p. 793–802, doi: 10.1144/gsjgs.141.5.0793. Polanski, J., 1970, Carbónico y Pérmico de la Argentina: Buenos Aires, Editorial Eudeba, p. 216. Ramos, V.A., 1985a, El Mesozoico de la Alta Cordillera de Mendoza: Facies y desarrollo estratigráfico, Argentina: IV Congreso Geológico Chileno, Actas 1, p. 492–513, Antofagasta.

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Ramos, V.A., 1985b, El Mesozoico de la Alta Cordillera de Mendoza: Reconstrucción tectónica de sus facies, Argentina: Antofagasta, IV Congreso Geológico Chileno, Actas 1, p. 104–118. Ramos, V.A., 1988a, The tectonics of the central Andes, 30° to 33°S latitude, in Clark, S.P., Jr., Burchfiel, B.C., and Suppe, J., eds., Processes in continental lithospheric deformation: Geological Society of America Special Paper 218, p. 31–54. Ramos, V.A., 1988b, Tectonics of the Late Proterozoic–Early Paleozoic: A collisional history of southern South America: Episodes, v. 11, p. 168–174. Ramos, V.A., 1989, The birth of southern South America: American Scientist, v. 77, p. 444–450. Ramos, V.A., 1993, Interpretación tectónica, in Ramos, V.A., ed., Geología y recursos naturales de Mendoza: Buenos Aires, XII Congreso Geológico Argentino y II Congreso de Exploración de Hidrocarburos (Mendoza), Relatorio, p. 257–266. Ramos, V.A., 1999, Plate tectonic setting of the Andean Cordillera: Episodes, v. 22, p. 183–190. Ramos, V.A., 2004, Cuyania, an exotic block to Gondwana: Review of a historical success and the present problems: Gondwana Research, v. 7, p. 1009–1026, doi: 10.1016/S1342-937X(05)71081-9. Ramos, V.A., and Basei, M., 1997, The basement of Chilenia: An exotic continental terrane to Gondwana during the Early Paleozoic: New Zealand, Symposium on Terrane Dynamics 9, p. 140–143. Ramos, V.A., and Kay, S.M., 1991, Triassic rifting and associated basalts in the Cuyo basin, central Argentina, in Harmon, R., and Rapela, C., eds., Andean magmatism and its tectonic setting: Geological Society of America Special Paper 265, p. 79–91. Ramos, V.A., and Yrigoyen, M.R., 1987, Geología de la región del Aconcagua, provincia de Mendoza: X Congreso Geológico Argentino, Actas, v. 4, p. 267–271. Ramos, V.A., Jordan, T.E., Allmendinger, R.W., Kay, S.M., Cortés, J.M., and Palma, M.A., 1984, Chilenia: Un terreno alóctono en la evolución paleozoica de los Andes Centrales: IX Congreso Geológico Argentino, Actas 2, p. 84–106. Ramos, V.A., Jordan, T.E., Allmendinger, R.W., Mpodozis, C., Kay, S.M., Cortés, J.M., and Palma, M.A., 1986, Paleozoic terranes of the central Argentine Chilean Andes: Tectonics, v. 5, p. 855–880. Ramos, V.A., Munizaga, F., and Kay, S.M., 1991, El magmatismo cenozoico a los 33°S de latitud: Geocronología y relaciones tectónicas: VI Congreso Geológico Chileno, Actas, v. I, p. 892–896. Ramos, V.A., Cegarra, M., and Cristallini, E., 1996a, Cenozoic tectonics of the High Andes of west-central Argentina, (30°–36°S latitude): Tectonophysics, v. 259, p. 185–200, doi: 10.1016/0040-1951(95)00064-X. Ramos, V.A., Aguirre Urreta, M.B., Álvarez, P.P., Cegarra, M., Cristallini, E.O., Kay, S.M., Lo Forte, G.L., Pereyra, F., and Pérez, D., 1996b, Geología de la región del Aconcagua, provincias de San Juan y Mendoza: Dirección Nacional del Servicio Geológico, Anales, v. 24, p. 1–510. Ramos, V.A., Cristallini, E., and Pérez, D.J., 2002, The Pampean flat-slab of the Central Andes: Journal of South American Earth Sciences, v. 15, no. 1, p. 59–78, doi: 10.1016/S0895-9811(02)00006-8. Rapalini, A.E., 1989, Estudio paleomagnético del vulcanismo permotriásico de la región andina de la República Argentina [Tesis doctoral]: Buenos Aires, Universidad de Buenos Aires, unpublished, 278 p. Rebori, L.O., 1979, Contribución al conocimiento geológico del extremo sur de la Precordilleramendocina: Sector Agua del Pizarro, Cacheuta, provincia de Mendoza: San Juan, Universidad Nacional de San Juan, Facultad de Ciencias Exactas y Naturales, Trabajo Final de Licenciatura (unpublished), 123 p. Schiller, W., 1912, La Alta Cordillera de San Juan y Mendoza y parte de la provincia de San Juan: Buenos Aires, Anales del Ministerio de Agricultura, Sección Geología, Mineralogía y Minería 7, no. 5, p. 1–68. Thomas, W.A., and Astini, R.A., 1999, Simple-shear conjugate rift margins of the Argentine Precordillera and the Ouachita embayment of Laurentia: Geological Society of America Bulletin, v. 111, p. 1069–1079, doi: 10.1130/0016-7606(1999)111<1069:SSCRMO>2.3.CO;2. Thomas, W.A., and Astini, R.A., 2003, Ordovician accretion of the Argentine Precordillera terrane to Gondwana: A review: Journal of South American Earth Sciences, v. 16, p. 67–79, doi: 10.1016/S0895-9811(03)00019-1. Thomas, W.A., Astini, R.A., and Bayona, G., 2002, Ordovician collision of the Argentine Precordillera with Gondwana independent of Laurentian Taconic orogeny: Tectonophysics, v. 345, p. 131–152, doi: 10.1016/ S0040-1951(01)00210-4.

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Triep, E.G., 1987, La falla activada durante el sismo principal de Mendoza de 1985 e implicancias tectónicas: X Congreso Geológico Argentino, Actas, v. 1, p. 199–202. Valencio, D.A., Vilas, J.F., and Pacca, I.G., 1983, The significance of the paleomagnetism of a sequence of red beds of the Middle and Upper sections of Paganzo Group (Argentina) and the correlation of Upper Paleozoic Lower Mesozoic rocks: Geophysical Journal of the Royal Astronomical Society, v. 51, p. 59–74. von Huene, R., Corvalán, J., Flueh, E.R., Hinz, K., Korstgard, J., Ranero, C.R., Weinrebe, W., and the Condor scientists, 1997, Tectonic control of the subducting Juan Fernandez Ridge on the Andean margin near Valparaíso, Chile: Tectonics, v. 16, p. 474–488, doi: 10.1029/96TC03703. Yañez, G., Ranero, G.R., von Huene, R., and Diaz, J., 2001, Magnetic anomaly interpretation across a segment of the southern Central Andes

(32–34°S): Implications on the role of the Juan Fernández ridge in the tectonic evolution of the margin during the Upper Tertiary: Journal of Geophysical Research, v. 106, no. B4, p. 6325–6345, doi: 10.1029/2000JB900337. Yrigoyen, M., 1976, Observaciones geológicas alrededor del Aconcagua: I Congreso Geológico Chileno, Actas 1, p. 169–190. Yrigoyen, M.R., 1979, Cordillera Principal, in Turner, J.C.M., ed., Segundo simposio de geología regional Argentina: Córdoba, Academia Nacional de Ciencias, v. 1, p. 651–694. Zeil, W., 1981, Volcanism and geodynamics at the turn of the Paleozoic to the Mesozoic in the Central and Southern Andes: Stuttgart, Zentralblatt für Geologie und Paläontologie, Teil I-1981(3/4), p. 298–318. MANUSCRIPT ACCEPTED BY THE SOCIETY 10 JANUARY 2008

Printed in the USA

The Geological Society of America Field Guide 13 2008

Field trip guide: Evolution of the Pampean flat-slab region over the shallowly subducting Nazca plate Víctor A. Ramos* Laboratorio de Tectónica Andina, Universidad de Buenos Aires, Consejo Nacional de Investigaciones Científicas y Técnicas (CONICET), Buenos Aires, Argentina ABSTRACT This field guide provides an opportunity to examine the central Andes between 31° and 32°S latitude in a segment characterized by flat-slab subduction. The field trip road was chosen to observe the westernmost contact between the basement uplift of Sierras Pampeanas and Precordillera, the early Paleozoic stratigraphy, and the Andean structure of the Precordillera, as well as a complete section of the Frontal and Principal Cordilleras in Argentina and Chile. The trip ends in the Coastal Cordillera along the Pacific margin. This road log discusses a complete early and late Paleozoic history of the central Andes with their typical Famatinian and Gondwanan orogenic rocks and the accretionary evolution of the Pacific margin at these latitudes. Superimposed on this framework, the structure of the Andes is viewed through the examination of the Precordillera and the Aconcagua fold-and-thrust belts, together with the observation of the Andean volcanic history, will allow reconstructing the shallowing of the subduction zone through the Neogene and the final formation of the Pampean flat-slab. Keywords: Andes, tectonics, terrane, allochthonous, Aconcagua. present volcanic activity, and therefore Late Cenozoic shortening is directly responsible for the present uplift and convergence rates. The route shows the different structural styles of the Precordillera, the Cordilleras Frontal and Principal along the Argentine slope, and the Cordilleras Principal and La Costa on the Chilean side, as indicated on the field trip road map (Fig. 1). The field trip ends in the Pacific coast.

INTRODUCTION This field trip provides the opportunity to examine the tectonic evolution of the central Andes in one of the most classic sections. The aim is to show the key localities where field data have been obtained and different regional relationships and tectonic models have been developed. The short duration and long distance to be covered does not allow a comprehensive review of the geology of the central Andes. The main objective of the trip is to view one of the most complete traverses of the Andes, where a noncollisional orogenic belt reaches elevations near 7 km (the highest mountains of the Western Hemisphere). These mountains are in an area with no

Brief Review of the History of Geologic Exploration and Research The geologic understanding of the central Andes started with the pioneering explorations of Charles Darwin in 1835, who was the first to describe the marine Mesozoic deposits deformed by faults (Fig. 2) along the present road that crosses the High Andes.

*[email protected]

Ramos, V.A., 2008, Field trip guide: Evolution of the Pampean flat-slab region over the shallowly subducting Nazca plate, in Kay, S.M., and Ramos, V.A., eds., Field trip guides to the Backbone of the Americas in the southern and central Andes: Ridge collision, shallow subduction, and plateau uplift: Geological Society of America Field Guide 13, p. 77–116, doi: 10.1130/2008.0013(04). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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The field trip provides the opportunity to examine the profile described by Darwin and the mountain shelters where Darwin stayed during his mountain crossing. Several German naturalists were sent by the Academia Nacional de Ciencias and the Museo de La Plata to geologically explore the High Andes of San Juan and Mendoza. The early observations of German Burmeister in 1857–1858, the structure outlined by Stelzner (1873, 1885), and the descriptions of Wehrli and Burckhardt in 1898 presented the concept that the Andes were a relatively simple mountain chain without the thrusts and overthrusts known in other mountain chains at that time. Later, in 1906 and 1907, Walter Schiller, a young geologist and mountain climber working in the Argentine Geological Survey conducted the first reconnaissance of the region (Fig. 3). As a result of his work, many structural complexities and important thrusting were recognized (Schiller, 1907, 1912). At the same time, the Precordillera was also explored, and its stratigraphy, structural geology, and mineral resources were described, making this area one of the best known regions of the

time (Stelzner, 1885; Bodenbender, 1902; Stappenbeck, 1910; Keidel, 1921; Bracaccini, 1946; etc.). Over the past several decades, many studies have provided a large volume of information on different aspects of the sedimentology, structure, and geologic evolution of this part of the central Andes (see Caminos, 2000, for a review). MAJOR GEOLOGICAL PROVINCES This segment of the central Andes has been divided into morphostructural units or geological provinces based on structural styles, geologic evolution, and morphological expression (Fig. 4). Sierras Pampeanas This geological province, located in central Argentina, is characterized by a series of crystalline basement blocks of Precambrian–Early Paleozoic age, which were uplifted and tilted

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Figure 2. Structural sketch section of Puente del Inca after Darwin (1846).

Figure 3. Classic section of Puente del Inca by Schiller (in Gerth, 1926).

during Tertiary Andean compression (González Bonorino, 1950a) in association with an episode of shallow subduction (Jordan et al., 1983a, 1983b; Kay et al., 1987; Ramos et al., 2002). The resulting structures closely resemble those of the Laramide region of the United States (Jordan and Allmendinger, 1986). The basement is composed of metamorphic and igneous rocks which correspond to two distinct orogenic cycles. The oldest Brasiliano or Pampean cycle is preserved along the eastern Sierras Pampeanas, and the metamorphic facies and igneous rocks define a north-south–trending belt of Late Proterozoic–Early Cambrian age (600–520 Ma) (Ramos, 1988a). The younger cycle is characterized by outcrops, along the western Sierras Pampeanas, which define the Famatinian orogen, an Early Paleozoic magmatic belt that reached its magmatic climax between 490 and 460 Ma (see reviews in Pankhurst and Rapela, 1998; Ramos, 2004). This metamorphic basement was partially covered by a series of continental deposits consisting of the Late Paleozoic Paganzo Group and Tertiary synorogenic deposits in alluvial and fluvial facies related to the uplift of the Sierras Pampeanas. Locally, some Triassic and Cretaceous continental sequences were deposited along rift basins developed in the eastern and western margins of the Sierras Pampeanas.

Precordillera The Precordillera is an Andean fold-and-thrust belt sequence with a typical thin-skinned structure that has developed in an early Paleozoic carbonate platform (Baldis and Chebli, 1969; see overview in Astini and Thomas, 1999). The western edge of the Precordillera coincides with a longitudinal depression known as the Iglesia-Calingasta-Uspallata valley. This tectonic trough is similar to the Canadian Rocky Mountain trench. In both regions, the modern morphology is controlled by the old Paleozoic continental margin (Price, 1981; Baldis et al., 1982). The Early Paleozoic history is represented by a carbonate platform of Early Cambrian to Middle Ordovician age. A fine biostratigraphic zonation has been defined in these highly fossiliferous deposits (Bordonaro, 1980; Baldis et al., 1982). Clastic marine Middle to Upper Ordovician rocks cover the platform in the eastern and central sectors, while slope and oceanic facies occur to the west. This Early Paleozoic continental margin existed until the Middle to Late Devonian. The marine sequences are mainly turbiditic facies, typical of flysch deposits (González Bonorino, 1975).

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Figure 4. Major morphostructural units of the central Andes (28°–33°S).

A major deformation, known as the Chanic, affected these Early Paleozoic rocks during the Middle to Late Devonian, developing a series of isolated ranges known as the “Protoprecordillera” (Amos and Rolleri, 1965). The previous slope facies were covered by continental Early Carboniferous alluvial deposits in the western Precordillera. Late Paleozoic rocks are deposited in a foreland basin, mainly littoral marine facies in the western Precordillera and continental fluvial to alluvial deposits in the central and eastern Precordillera. Continental redbeds of the Paganzo Group rocks overlap the eastern Precordillera. In well-known localities such as Barreal and Rinconada among others, excellent outcrops of the Gondwana glacial deposits are preserved in both marine and continental facies (Keidel, 1921; Du Toit, 1927; López Gamundi and Amos, 1986). Synorogenic Tertiary deposits permit the reconstruction of the Precordillera thrust sequence, which began ca. 18 Ma in the northwestern side and continues until the present day on the eastern side (Jordan et al., 1997). Cordillera Frontal The Cordillera Frontal is composed of units that formed during the Gondwanides orogeny in the Late Paleozoic to early Mesozoic. These units result from Andean-type subduction,

followed by generalized extension. Most of the rocks of this province are Late Paleozoic–Triassic andesitic to silicic magmatic rocks of the Choiyoi Group (Caminos, 1979). During the Andean deformation, the Cordillera Frontal behaved as a rigid block, as shown by the presence of thick-skinned thrusts (Polanski, 1957, 1970). Scattered exposures of the pre–Late Paleozoic basement of this province occur in Argentina and Chile. They consist of low- to medium-grade metamorphic rocks of latest Proterozoic to Early Cambrian age. The Early Paleozoic is represented by isolated outcrops of Silurian-Devonian marine limestones which are covered by widespread Carboniferous–Early Permian turbidites. This Late Paleozoic facies contrasts with the littoral facies of Precordillera to the east (Caminos, 1979). Volcanic activity started in the Early to Middle Carboniferous with subduction related andesites, dacites, and rhyolites. A subsequent period of generalized extension from Middle Permian up Early Triassic times resulted in the thick pile of Choiyoi rhyolites and associated granites (Kay et al., 1989; Llambías and Sato, 1990). These volcanic rocks, which reach thickness of up to 2–4 km in the Cordillera del Tigre, unconformably overlie the older rocks. Deformation of the Carboniferous–Early Permian rocks occurred in the middle Permian San Rafael orogenic phase (Ramos, 1988a; Rapalini, 1989).

Evolution of the Pampean flat-slab region The boundary between the Cordillera Frontal and the Precordillera was the locus of Triassic rifting that is associated with up to 2 km of synrift deposits, scattered alkaline basalts and associated sag facies. The rift developed in the hanging wall of the suture between Cuyania and Chilenia terranes (Ramos and Kay, 1991). The early sedimentation of the rift was coeval with the Choiyoi volcanism. Cordillera Principal The Cordillera Principal, or Main Andes, was the locus of the Andean orogeny during latest Mesozoic and Cenozoic times. Jurassic and Cretaceous marine deposits were deformed in different styles depending on the extent of participation of the basement in the deformation. In the northern sector, a thick-skinned tectonic style is described by Moscoso and Mpodozis (1988), whereas to the south thin-skinned structures such as in the Aconcagua fold-andthrust belt developed (Yrigoyen, 1976, 1979; Ramos, 1988b; Ramos et al., 1996a, 1996b). A thick sequence of marine Mesozoic deposits unconformably overlies the Carboniferous flysch and the Choiyoi volcanics of the Cordillera Frontal. Several sedimentary cycles are recognized from the Early Jurassic to the Early Cretaceous (Groeber, 1946; Legarreta and Gulisano, 1989). These cycles begin with black shales, sandstones, and limestones and terminate with thick gypsum levels and continental red beds. Abundant ammonites permitted a biostratigraphic zonation of these deposits (Riccardi, 1984). Along the continental divide these sedimentary sequences interfinger with volcanic and pyroclastic rocks of Late Jurassic– Early Cretaceous age. The volcanic pile, which can be up to 6 km thick in the Chilean side, has a burial metamorphism typical of that developed in a high thermal gradient during active subsidence (Levi et al., 1982). These volcanic sequences occur along a western inner arc developed between the Cordilleras de la Costa and Principal, and an eastern outer arc along the present international border. An intra-arc basin between the two arcs is filled with shallow marine and continental deposits (Charrier, 1973; Charrier et al., 2005; Rivano et al., 1985; Ramos, 1985b). Most of the early-middle Mesozoic was dominated by an extensional regime closely linked with the early stages of the opening of the South Atlantic (Uliana et al., 1989). The beginning of the drift phase in the Atlantic during the middle Cretaceous changed the tectonic regime to the present Andean compressional stage. A series of volcanic arcs shifted from the Cordillera de la Costa in the Jurassic to the Cordillera Principal in the Late Tertiary (Ramos, 1988b). The volcanic and volcaniclastic rocks interbedded with alluvial-fan facies. Glacial deposits from four different glaciations are widespread in the main valleys, representing alpine type glaciations during Pliocene and Quaternary times.

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Cordillera de la Costa Along the present continental margin, Late Paleozoic metamorphic rocks are preserved that represent pieces of an accretionary prism developed in the Late Paleozoic (Hervé, 1988). Emplaced in this metamorphic basement are a series of magmatic belts of Jurassic and Cretaceous age. Most of this region is suspected to have significant latitudinal motion (Forsythe et al., 1986; Mpodozis and Ramos, 1990). The generalized stratigraphy of the Cordillera de la Costa is as follows. On the eastern flank of the Cordillera de la Costa, Mesozoic marine sequences that developed west of the Mesozoic volcanic rocks are preserved. These marine deposits commonly interfinger with volcanic rocks that can be associated with Cu manto-type deposits. Isolated patches of accretionary prism deposits, which are principally Paleozoic to early Mesozoic in age, crop out along the Pacific coast. MODERN PLATE TECTONIC SETTING This segment of the central Andes between 28° and 33°S has a distinctive plate tectonic setting. The present convergence rate between the subducted Nazca plate and the South American plate averages ~9 cm per year. Earthquake locations delineate a Benioff zone that is gently dipping to the east defining a shallow subduction zone (see Fig. 5) (Cahill and Isacks, 1992; Pardo et al., 2002). This flat subduction segment is characterized by an almost flat section at ~100 km dept that is flanked to the north and south by steeper segments that dip ~30° eastward. A corresponding tectonic segmentation exists in the plate above the Benioff zone. The most obvious and consistent correlation is between Quaternary volcanism and the dip of the subducted slab. Quaternary volcanism is absent in the subhorizontal segment. The development of the Sierras Pampeanas geological province is controlled by flat subduction. Present tectonic shortening is principally concentrated along a narrow belt between this province and the Precordillera. Global positioning system (GPS) displacements analyzed by Brooks et al. (2003) show a strong gradient between both provinces. Intraplate earthquake concentrations have been found in the basement of the eastern Precordillera and the western Sierras Pampeanas, in close coincidence with the superficial neotectonic activity. Focal mechanisms indicate east-west contraction with null to minor strike-slip displacements (Chinn and Isacks, 1983; Pardo et al., 2002). The origin of this flat subduction segment has been attributed to the approach and collision of aseismic ridges to the Pacific continental margin (Pilger, 1981; Yañez et al., 2001), to changes in the age of the subducted oceanic crust (Wortel, 1984), and to the differential shortening of a previously weakened hot continental crust (Isacks, 1988; Cahill and Isacks, 1992). See recent review in Ramos et al. (2002).

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Figure 5. Sections showing earthquakes hypocenters (A) north and in (B) the flat subduction segment (after Cahill and Isacks, 1992). Vn—volcano.

TERRANE HISTORY The existence of oceanic rocks separating the Cordillera Frontal from the Precordillera has attracted the attention of geologists since the early work of Borrello (1969). These oceanic rocks have been interpreted as indicating a suture between different continental terranes (Ramos et al., 1984, 1986). Consequently, several other sutures have been identified (Ramos, 1984, 1988a, 1988b; Mpodozis and Ramos, 1990; Astini et al., 1995, 1996, 1999). A map of the suggested terranes in the region after Ramos (1988a) is shown in Figure 6. The first attempt to explain the presence of olenellid trilobites in South America was made by Ross (1975), who explained

them by larval transfer by oceanic currents. The location of this fauna with olenellid trilobites outside of Laurentia was intriguing, mainly because this fauna was only known in the ancestral North American craton and in the northwestern British Isles. Several years later, the striking coincidence in the subsidence curves of the Appalachians carbonates and the Precordillera carbonate platform, induced Bond et al. (1984) to suggest that both regions were conjugate margins and shared a common rift drift transition. The shared carbonate platformal history and faunal provinciality between the Precordillera and the Laurentian margin of the Appalachians led Ramos et al. (1986) to propose that the Precordillera was a far-traveled terrane derived from the northern Appalachians. This proposal was refined by Mpodozis

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Figure 6. Proposed sutures and allochthonous terranes in the central Andes (after Ramos, 1988b).

and Ramos (1990), and Astini et al. (1995, 1996), within the tectonic framework of the early Paleozoic basement of the Andes. Two different models have been proposed to explain the accretion of the Precordillera to the protomargin of Gondwana. One of the models proposed an independent microcontinent or microplate, detached from Laurentia during Early Cambrian time, which collided against Gondwana during Middle to Late Ordovician time (Ramos et al., 1986; Benedetto and Astini, 1993; Astini et al., 1995, 1996). The other model proposed a continent to continent collision between Laurentia and Gondwana during Early to Middle Ordovician time from 487 to 467 Ma. Subsequently, during the Late Ordovician, the separation of the two continents left behind the Precordillera terrane on the Gondwanan side, with the opening of an ocean on the western side of Precordillera (Dalla Salda et al., 1992a, 1992b; Dalziel, 1992, 1993, 1997; Dalziel et al., 1994, 1996). Both proposals required an active early Paleozoic margin in the Sierras Pampeanas, and explained the magmatism and the Ordovician deformation known as the Ocloyic event as the result of a collisional orogeny (Ramos et al., 1986; Dalla Salda et al., 1992a, 1992b). The microcontinent hypothesis required a second early Paleozoic accretion to close the ocean that bounded the western Precordillera. The accretion of the Chilenia microcontinent produced a second foreland basin and a shifting of the magmatic activity to the Pacific margin (Ramos et al., 1984). This colli-

sion occurred either in Late Devonian time (Ramos et al., 1986) or as proposed by Astini (1996) in Early Devonian time. Sedimentological studies of these foreland basins and geochronologic data of peak metamorphism and associated deformation suggest an Early Devonian age for the beginning of the accretion of Chilenia. The challenging pre-Pangea Southwest United States– East Antarctic (SWEAT) reconstruction of the late Proterozoic continents proposed by Moores (1991), complemented by Dalziel’s 1991 Laurentian end run, may have provided a mechanism for the transfer of the Precordillera. See recent reviews by Thomas and Astini (1999, 2003) and Ramos (2004). Pampia Terrane The basement of the central and western Sierras Pampeanas constitute an independent terrane that was accreted to the Río de La Plata Craton during the Late Proterozoic–Early Cambrian times (Ramos, 1988a; Kraemer et al., 1995). A subduction zone dipping toward the Río de la Plata craton was responsible for the development of a magmatic arc, which is represented by a series of gabbros, tonalites, and granitoids in the eastern Sierras Pampeanas (Lira et al., 1997). These rocks have ages ranging between 700 and 525 Ma (Rapela et al., 1998). Final amalgamation and an uplift of 15 km (Ramos, 1988b) occurred during the Early Cambrian (Rapela et al., 1998). The magmatic rocks exposed along

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the western border have petrological and geochemical characteristics typical of an active continental margin (Rapela et al., 1992; Pankhurst et al., 1998; Quenardelle and Ramos, 1999). The deformational history and the age of the magmatism are consistent with the western Sierras Pampeanas representing an active margin from the Late Cambrian until the Middle to Late Ordovician (Ramos, 1989, 1993; Pankhurst et al., 1998). The western border of the Sierras Pampeanas is also characterized by a synthetic fold-andthrust belt with west vergence (Dalla Salda, 1987). Cuyania Composite Terrane Petrologic studies and dating performed in the basement of Precordillera and adjacent terranes (Kay et al., 1996; Sato et al., 2000) have demonstrated that Precordillera was part of a composite terrane of Grenvillian age. This terrane, named Cuyania, was accreted as a whole to the Gondwana margin (Ramos et al., 1996b; Thomas and Astini 1996, 1999, 2003; Astini and Thomas, 1999). A map of the composite Cuyania terrane is shown in Figure 7. In addition to the olenellid trilobites, other faunal data from Cambrian and Ordovician rocks of the Precordillera terrane were examined by Benedetto et al. (1999) to refine their biogeographic relationships with Laurentia and Gondwana. The study, based on benthic organisms such as sponges, bryozoans, brachiopods, bivalves, and ostracods, led to the recognition of four successive stages of Precordilleran biogeographic evolution: (1) Laurentian stage (Cambrian-Tremadoc), (2) isolation stage (Arenig–early Llanvirn), (3) pre-accretion stage (Llanvirn-Caradoc), and (4) Gondwanan stage (Hirnantian-Silurian). Each stage denotes its specific position during the rifting-drifting-collision sequence and each one reflects a different pattern of faunal exchange. During the Laurentian stage, the nearly complete identity with Appalachian faunas supports a close geographic connection between Precordillera and Laurentia. The isolation stage begins when taxa that have not been recorded in Laurentia appear for the first time in the Precordillera basin. Through this stage, the Laurentian faunal influence decreases and a number of endemic Baltic-Avalonian genera correlatively increase. The pre-accretion stage is characterized by a paucity of Laurentian forms, the arrival of Gondwanan taxa, and an unusually high level of endemicity. The latter may reflect a degree of geographic isolation, and may in part be due to biologic factors related to dispersal mechanisms. The Gondwanan stage starts after the accretion of the Cuyania terrane at the end of the Ordovician. The review of faunal data supports the “far traveled microplate” hypothesis which is generally consistent with the geological evidence. The Precordillera, a part of the Cuyania composite terrane, was interpreted by Astini et al. (1995, 1996), Thomas and Astini (1996, 1999, 2003), and Astini and Thomas (1999), as a fragment of rifted Laurentian continental crust and passive-margin cover. Two separate episodes of extension, ~60–70 m.y. apart, and a contractional event are recorded in the early Paleozoic history of the Precordillera. Crustal extension during the Early Cam-

brian (possibly starting in the latest Proterozoic) led to asymmetric continental rifting and separation of Precordillera from the Ouachita embayment of southern Laurentia (Thomas and Astini, 1996). Synrift graben-fill successions of the Precordillera are overstepped by latest Early Cambrian carbonates indicating riftto-drift transition and initiation of passive margin deposition. Faunal evolution in the uppermost Lower Cambrian through Lower Ordovician passive margin succession suggests isolated drifting of the Precordillera as a Laurentian orphan across the Iapetus Ocean from the Late Cambrian to the Early Ordovician. Subsidence curves are typical of post-rift thermal subsidence on rifted continental margins (see Thomas and Astini, 1999, 2003). A contractional event, interpreted as the docking of the Precordillera with Gondwana, is documented by mylonitic fabrics, Ocloyic metamorphic ages (464 Ma) imprinted on Grenville basement rocks (Ramos et al., 1998), and west-directed thrusting of passive-margin limestones in eastern Precordillera. The collision is confined to early Middle Ordovician time. A second extensional episode, during the Middle Ordovician starting in the Llanvirn, is evidenced by irregular distribution of sediments and hiatuses, abrupt changes in lithofacies, and local slope-scarp facies associated with block faults. This extension may have been related to pre-collisional flexural extension (Astini et al., 1996). Paleomagnetic data on the Early Cambrian rocks of the Precordillera (Rapalini et al., 1999) indicate paleolatitudes of approximately 20°, similar to the Ouachita embayment of Laurentia, as inferred by Thomas and Astini (1996, 1999, 2003). Pie de Palo is the other terrane that constitutes the Cuyania composite terrane. The Precambrian Grenvillian basement of this terrane is composed of an ophiolitic assemblage (Vujovich and Kay, 1998) of Middle Proterozoic age, and a series of gneisses and amphibolites of similar age. U/Pb ages in zircons, as well as Ar/Ar ages, have demonstrated that this basement was already part of the same terrane as the Precordillera when it collided against the protomargin of Gondwana during Middle-Late Ordovician times (Ramos et al., 1998). Several deformation phases have been recorded in the eastern border of Precordillera during the Late Ordovician and Silurian (Baldis et al., 1984). Particularly spectacular are the tightly deformed flysch deposits of the Late Ordovician–Silurian Rinconada Formation that contain olistoliths of Early Ordovician rocks. Most of this deformation lasted until the OrdovicianSilurian boundary, at the time when the Precordillera was in the final stages of amalgamation with the Sierras Pampeanas. The Cambro-Ordovician magmatic belt of the western Sierras Pampeanas also records a series of episodes that can be matched to the sedimentary evolution of the Early Paleozoic sequences of Precordillera. A coeval evolution can be integrated in a single simple model (Ramos, 2004). The deformation recorded in the sedimentary sequences also matches the metamorphic episodes that are indicated by 40Ar/39Ar ages at Pie de Palo (Ramos et al., 1998). The western border of the Precordillera was the continental margin during most of the Early Paleozoic, as shown by the sedimentary facies in the Sierra de Tontal (Cingolani et al., 1989).

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The reconstruction of this continental margin is based on sedimentologic and paleontologic evidence (Baldis et al., 1982), as well as geochemical characteristics of the ophiolitic assemblages developed along the western border of Precordillera (Kay et al., 1984; Haller and Ramos, 1984; Davis et al., 1999). Imbricated structures in the Ordovician and Siluro-Devonian rocks indicate a strong deformation during the Middle to Late Devonian. The western vergence of these structures can be seen in the Sierras de Cortaderas and Sandalio (Cortés, 1989). This deformation, known as the Chanic event, has been interpreted as the result of the collision of the Chilenia terrane against the western Precordillera continental margin (Ramos et al., 1984, 1986). Chilenia Terrane The basement of this large terrane underlies the Main Andes of Argentina and Chile, and is exposed in a series of minor erosional windows or is preserved as roof pendants in some of the large granitic batholiths. The basement is known in the Cordón del Plata and Cordón del Portillo (Caminos, 1965), and in the La Pampa gneisses (Mpodozis and Ramos, 1990). The metamorphic rocks have been dated at Las Yaretas in the Cordillera Frontal where U/Pb ages in zircon indicate 1069 ± 36 Ma (Ramos and Basei, 1997). This age, combined with the absence of Brasiliano deformation, suggests Laurentian affinities for the Chilenia terrane. The large amount of upper Paleozoic granitoid on both slopes of the Andes in Argentina and Chile has been also taken as evidence for the continental nature of the basement by Nasi et al. (1985). Isotopic (initial 87Sr/86Sr ratios) and geochemical characteristics of these magmatic rocks are consistent with a crustal component of Precambrian age in the source of these magmas (Mpodozis and Kay, 1990). The Early Paleozoic cover is scarce and different from Precordillera. Accretion to Gondwana is postulated during Middle-Late Devonian times (Ramos et al., 1986). Final amalgamation occurred in Early Carboniferous times, where a continental to shallow marine deposits overstepped both terranes. Carboniferous marine units were the first sedimentary rocks amalgamated to the Chilenia terrane. Continental deposits are widespread in the Sierras Pampeanas and Eastern Precordillera; nearshore marine facies occur in the Western Precordillera and estuarine to turbiditic facies are present along the Cordilleras Frontal and Principal. An important subduction-related magmatic activity is recorded during the Late Paleozoic. This magmatic belt, when compared with the location of the Early Paleozoic magmatic belt, was shifted more than 300 km toward the ocean (Ramos et al., 1986). These entire Carboniferous to Early Permian units were strongly deformed during the San Rafael orogenic phase (Ramos, 1988b). Some authors have attributed the cessation of subductionrelated magmatism and the consequent middle Permian deformation to the collision of an unidentified terrane (terrane Equis of Mpodozis and Kay, 1990). Other authors interpreted the cessation of magmatism as being related to and episode of flat subduction during the middle Permian that produced the strong San Rafael deformation (Martínez et al., 2006).

Soon after the San Rafael deformation, a generalized extension took place associated with a period in which the Gondwana plate was stationary with respect to the South Pole, as shown by the polar wandering path (Valencio et al., 1983; Ramos, 1988b). Magmatism was widespread during this extensional period (Zeil, 1981), and as a consequence, batholiths like the Colangüil were emplaced between 264 and 247 Ma (Llambías and Sato, 1990). Magmatic activity in this period is represented by rhyolitic volcanism within the Paleozoic accreted terranes, which are presently exposed in the Cordillera Frontal, western Sierras Pampeanas, and Precordillera. This felsic volcanism known as the Choiyoi province (Kay et al., 1989) was interpreted as evidence of generalized extension during Triassic times (Zeil, 1981) and post-collisional volcanism associated with slab breakoff in the final stages of the amalgamation of Pangea (Mpodozis and Kay, 1992). These rhyolitic rocks have alternatively been related to the steepening of a shallow subduction zone starting in the Late Permian (Martínez et al., 2006). ANDEAN TECTONICS The Cenozoic sedimentary history records the eastward migration of the orogenic front (see Fig. 8). Thick sequences of continental deposits (Santa María Conglomerates of Schiller, 1912) unconformably overlie the Mesozoic rocks in the Cordillera Principal. The angular unconformity is clearly seen east of Cerro Aconcagua and west of Cerro Penitentes. Those conglomeratic deposits are interpreted as alluvial fan sediments interfingered with the volcanics of the Farellones Formation (25–10 Ma, Munizaga and Vicente, 1982). A minimum age of 8.6 Ma was obtained in the continental deposits based on K/Ar dating of pyroclastic rocks interbedded in the uppermost section of the Santa María Conglomerates (Ramos et al., 1996b). The Tertiary deposits farther east of the High Cordillera are represented by distal fluvial facies partially synchronous with the Santa María Conglomerates and the volcanism of the Farellones Formation. The foreland Tertiary Andean sequences at these latitudes (30°–33°S) contain several tuff layers, which attest to cordilleran volcanic activity at that time. An unconformity separates La Pilona beds exposed in the Uspallata valley and the Cacheuta area from older Tertiary beds. This unconformity was produced in the late Miocene prior to the 10–8 Ma age tuffs (Ramos et al., 1996a, 1996b), and it was also coeval with the uplift of the Cordillera Frontal. The Late Miocene and Pliocene deposits of the Uspallata and Cacheuta regions were folded and thrust at that time with the subsequent deposition of the alluvial fan deposits of the Mogotes Formation during the Plio-Pleistocene. Therefore, the Tertiary sedimentary facies show a migration of the coarse alluvial fan facies from: (a) the inner area of Cerro Penitentes in the High Andes between 20 and 10 Ma to (b) the Uspallata valley and Cacheuta between 10 and 5 Ma, to (c) the outer foothills of the city of Mendoza between 2 Ma and the present active front. Even the Plio-Pleistocene fanglomerates of the Mogotes Formation cropping out west of the city of Mendoza

Figure 8. Nazca plate segmentation and main geologic setting of the Andes of Argentina and Chile (based on Jordan et al., 1983b).

Ramos

(Cerro de la Gloria) and other younger alluvial fans have been deformed by neotectonic activity in the Mendoza region. The seismic sections of the plains located eastward of the Precordillera clearly show that the present orogenic front is composed of a set of imbricated thrusts. This structural style was corroborated by drilling, and has somewhat similar characteristics to the previous fronts (Fig. 9). The thrust front in the eastern side of the Precordillera is still active. Intense compressive deformation, as seen in Sierra de las Peñas (see Cortés, 1990) and, as inferred from earthquake focal mechanisms and escarpments on the alluvial fans, is continuing today. The Andean structure of the central Andes is the result of a combination of several tectonic mechanisms. There is a striking coincidence between the increase of plate motion rates, the cessation of magmatism, and the compressive deformation at the orogenic fronts. The Andean history begins in the late Oligocene (Ramos, 1999). Most of the Oligocene was quiescent with localized extension (Godoy et al., 1999) coincident with the volcanic rift along the axis of the Cordillera. This extension ended at ca. 20 Ma when the volcanic activity of the Farellones Formation started (Munizaga and Vicente, 1982). Two interrelated tectonic features demonstrate that a change in the geometry of the Benioff zone followed: (a) eastward migration of the subduction-related magmatic centers from a position ~180 km from the present trench at ca. 25 Ma, to 700 km away from the trench at 2 Ma at the Central Sierras Pampeanas in the latest Pliocene (Ramos et al., 1991; Kay et al., 1987, 1991; Kay and Mpodozis, 2002); (b) geochemical characteristics indicate a thickening of the continental crust, between 18 Ma and the present, related to the tectonic stacking of the Andean Cordillera (Kay et al., 1987; 1991); and (c) eastward shifting of the orogenic deformation during the past 20 Ma: 275 km from the trench at 20–10 Ma, 325 km at 10–5 Ma, and 365 km at 2 Ma. This implies an average propagation rate of 2.5 mm/yr of the orogenic front in the past 20 m.y., although shifting was probably episodic. The causes of the change in the Benioff zone geometry and segmentation are probably complex and multifaceted: (a) The breakup of the Farallon plate into the Cocos and Nazca plates, which occurred at 25 Ma, seems to mark the beginning of a period of higher convergence rates (Handschumacher, 1976); this age coincides with the initiation of Farellones magmatism and is a milestone in the geodynamic evolution of the area; (b) the increase in plate convergence from 25 to 26 Ma up to 10 Ma as defined by Pilger (1984) and Pardo Casas and Molnar (1987) when the present deceleration began; and (c) several authors have explained the present segmentation of the subducted Nazca plate as being controlled by the collision of aseismic ridges (Pilger, 1981). In Pilger’s interpretation, the buoyancy effects produced by subduction of those ridges, combined with the younger age of the subducted slab, contribute to diminish the angle of the Benioff zone. The effects of the subduction of the Juan Fernández hot spot trace began ca. 15 Ma according to Pilger (1984). At the present latitudes evidence of shallowing of the subduction zone shows a similar trend of southward

Figure 9. Conceptual model of the crustal section of the Andes of Argentina and Chile based on gravimetric data and a limited deep seismic reflection data (after Ramos et al., 2004). F.T.B.—fold thrust belt.

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Evolution of the Pampean flat-slab region migration, which can be correlated with the southward shift of the Juan Fernández ridge collision along the trench (von Huene et al., 1997; Yañez et al., 2001). South of the region of flat subduction, there is an age decrease in the subducted slab and even with a higher thermal gradient there than in the flat-slab segment, the present angle of subduction is ~30° (Isacks et al., 1982). Isacks’ model (1988) indicates that the present dip of the different slab segments is controlled by the width of previous weakened zone of the upper plate. This zone is related to the size of the asthenospheric wedge between the oceanic and the continental plates. A greater amount of shortening occurred in the central segment because the weakened area was the widest, as indicated by the extension of the magmatic activity in the Puna Altiplano. In the analyzed segment (30°–31°S), a relatively narrow weakened zone produced the overriding of the Nazca plate by the South American plate associated with relatively minor shortening.

72°W

89

FIELD TRIP LOG The road map of Figure 10 shows the location of detail geologic maps of the different stops of the fieldtrip guide. This map should be complemented with a road map such as the Atlas Vial of the Automovil Club Argentino or similar commercial maps. DAY 1—ACTIVE TECTONICS OF EASTERN PRECORDILLERA From Mendoza to San Juan, Argentina The present setting of the Precordillera within the Pampean flat-slab segment of the central Andes records an intensive shortening along the eastern foothills (Fig. 11). Although the geometry of the cover is characterized by thin-skinned thrusting, important shortening is recorded in the middle-upper crust, where the basement is deformed as indicated by earthquakes at ~15 km depth.

30°S

70°W

71°W

FRO N DILL

ERA

Fig. 21

Fig. 18 y2 Da

Fig. 22 Fig. 17

SAN JUAN Fig. 15 Barreal

ORD IL

32°S

PRE C

COR

Fig. 14

Fig. 24

Aconcagua

Fig. 32 San Felipe

Uspallata

(6967m)

Fig. 12

Fig. 11 4 ay

MENDOZA

án uy

VALLE CENTRAL

za

Vn.Tupungato (6800m)

do en oM

VALPARAISO

Rí

Puente D del Fig. 25 Inca

Los Andes Fig. 30

oT un

Viña del Mar

(6350m)

Rí

Day 5

Day 3

(6769m) Ramada

DILL

ERA

Mercedario

Fig. 34

Río San Juan

LER

Fig. 23

A

PRIN

CIPA

L

COR

Fig. 16

Day 1

INA NT

ILE

GE

CH

AR

CORDILLERA DE LA COSTA

Pacific Ocean 32°S

TAL

40

7

SANTIAGO

50 km 71°W

70°W

34°S

Figure 10. Field trip road map with location of the detailed maps in this guide.

90

Ramos

Figure 11. Thrust front and Quaternary terraces west of Cerro de la Gloria and the city of Mendoza (based on Rodríguez and Barton, 1990). C-Or—Cambrian and Ordovician limestones; S-Dv—Silurian and Devonian sandstones; Tr—Triassic deposits; Tp—Divisadero Largo Formation; Tm—Mariño Formation; TQm—Mogotes Formation; Qt—Terraced deposits.

The migration of the thrust front reached the eastern Precordillera during late Pliocene times, as shown by the proximal facies of the synorogenic deposits of Los Mogotes Formation. Two thrust systems are seen in the segment considered here (see Fig. 12). The western system is characterized by a series of imbricated thrusts with east vergence detached in Early Paleozoic rocks, as shown in the Cerro La Cal, Divisadero Largo (Fig. 11), and Sierra de las Peñas sections. The thrust sheets are placing Triassic red beds over the distal and proximal facies of the Cenozoic strata. This system records an important seismotectonic activity as observed east of Cerro La Cal, in Río de las Peñas, and Cerro de la Gloria. The eastern system has west vergence at these latitudes and involves the Proterozoic–Early Paleozoic basement of Sierras Pampeanas. Miocene deposits are overridden by these metamorphic rocks. Farther north in the San Juan area, the present shortening is recorded by backthrusts with west vergence belonging to this eastern system (see Day 2). Stop 1-1: Cerro de la Gloria Cerro de la Gloria is a small hill located along the western border of the city of Mendoza (Fig. 11). Important earthquakes destroyed the city of Mendoza in 1861 and partially destroyed it in 1920 and 1985. The first earthquake is related to small

displacements on the Cerro La Cal thrust and had an estimated magnitude of Ms = 7.0 (Mingorance, 2004). The fault scarp can be seen across Las Heras Avenue, a few meters west from where it intersects Perú Street. The railroad seen here was built along the fault scarp a few years after the 1861 earthquake that destroyed the city of Mendoza. The 1985 earthquake had a hypocenter located a few kilometers southwest of the city (Ms = 5.75, depth 14 km), with a compressive focal mechanism. The most probable active fault plane was a reverse fault dipping 56° to the west (Triep, 1987). A series of east-verging thrusts imbricate Paleozoic rocks (C—Cambrian, Or—Ordovician, and Dv—Devonian), Triassic rift deposits (Tr), and Tertiary synorogenic deposits (Tp— Oligocene, Tm—Miocene, and TQm—Late Pliocene) with (Qt) Pleistocene alluvial fan deposits (see Fig. 11). The dominant north-south trend of the different thrusts along tens of kilometers is abruptly modified by an east-west deflection near the city of Mendoza. This deflection corresponds to a lateral ramp that is a reactivation of an east-west transfer fault of the Triassic rift system. North of this transfer fault, the main rift depocenter is to the east whereas to the south, the main depocenter is in the western side of the rift basin. The flat-lying to gently-dipping Pleistocene deposits (0° to 25°E) seen along the road show evidence of progressive unconformities. As we move to the west, dips steepen until reaching

Evolution of the Pampean flat-slab region

91

A

B Mendoza city

ria

a

G

t

lo

.L

Co

E

La

Divisadero Largo Fault

ul

Ce rro

CGx-1

Fa

Ca lF au lt

Seismic line control

W

1000 Tc Sea level Trsyn

Trsag t

a

meters

–1000 Ch –2000

–3000

Dv

–4000

El

Ch

a all

o

ult Fa

–5000 1 km –6000

Figure 12. Structural cross section of the foothills of the city of Mendoza. Location and references in Figure 11 (modified from Parke, 2001).

subvertical in the conglomerates of the Mogotes Formation (TQm—Late Pliocene). These sequences indicate the strong neotectonic activity of the thrust front during the Quaternary. The coarse conglomerates correspond to mass wasting flow and proximal alluvial fan facies. From the summit of Cerro de la Gloria, there is a general view of the thrust front and the different Quaternary terraces related to the Precordillera uplift. Some of the terraces are cut by younger Quaternary faults. The general geometry of the faults is controlled by the tectonic inversion of the Triassic rift system. A subvolcanic dacitic body of Tertiary age (Tv in Fig. 11) is one of the easternmost evidences of the Oligocene–Early Miocene volcanic arc at these latitudes. Stop 1-2: Borbollón Anticline The next stop is reached by taking Route 40 north just past the Mendoza airport (Fig. 13). The orogenic front in this region is formed by two actively growing structures, the Borbollón and Capdeville anticlines, bounded to the west by the Cerro La Cal Fault. Both structures have an east vergence. The hypocenter of the 1861 Mendoza earthquake was located along this fault in the segment north of Cerro La Cal. The Mendoza airport, located in the northern end of the city is sited along the axis of an actively growing anticline (Fig. 13). Seismic sections across the Borbollón anticline show progressive unconformities since at least late Miocene times (ca. 6 Ma, Olgiati, 2002). 40Ar/39Ar ages of 27.9 ± 0.6 k.y. for the white

ash-fall tuffs in the Borbollón anticline and 16.2 ± 0.6 k.y. in the Capdeville anticline confirm a Late Quaternary age. Both ashfall tuffs are derived from the active volcanic arc located to the southwest, south of the Tupungato volcano. At this stop, the western exposed limb of the anticline is dipping at ~3°W in contrast with late Miocene deposits that dip more than 25° at depth on the seismic line. Drive to the City of San Juan The drive is to the north from the city of Mendoza along National Highway 40 for ~150 km. The road goes along the boundary between distal alluvial fan and playa-lake deposits. On the western side of the road, the Sierra de las Peñas can be seen. This is one of the most active uplifting areas in Quaternary times in the region. Synorogenic deposits are cut by several fold scarps next to the southern plunge of the main structure and several progressive unconformities occur. Las Peñas thrust is one of the most spectacular fault scarps with more than 60 m of throw on Quaternary gravels. The alluvial fan deposits show several active fault scarps parallel to the mountain front. North of Río de las Peñas, a west-verging system intersects the mountain front. This west-verging system is a thick-skinned belt controlled by deep basement faults as depicted in the seismic lines and is characteristic of the adjacent Sierras Pampeanas. The minor uplifts could represent the still buried westernmost basement blocks of the Sierras Pampeanas.

92

Ramos 68°50′W 32°40′S

68°40′W

OROGENIC FRONT

7

32°40′S

Capdeville

Anticline

16.2 ± 0.6 k.y.

COl Cerro La Cal Qt

40

Estancia El Jagüé

TQ

Pzi

TQ

Borbolloncito

Borbo

Cerro La Cal

llón A

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nticlin e

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Legend

STOP 1-2 Borbollón 7

27.9 ± 0.6 k.y.

Qt

Terraced deposits

TQ

Borbollón beds

Pzi

Devonian sandstones

COl

Cambro-Ordovician 32°50′S limestones

Aeropuerto 32°50′S Pzi

Thrust Active thrust

Mendoza city

0

1

2 km 68°40′W

Figure 13. Geologic map of the orogenic front immediately north of the city of Mendoza. Stars show the proposed location of the 1861 Mendoza earthquake.

Stop 1-3: Cerro Valdivia Stop 1-3 is along Highway 40 a few kilometers before arriving in San Juan (see Fig. 10). This stop shows the Grenville-age basement rocks of the exotic Cuyania block that collided against Gondwana in Late Ordovician times. An U-Pb age in this sector yielded ages ca. 1100 Ma. The protoliths of these metamorphic rocks are island-arc deposits and a sedimentary cover similar to that in El Llano uplift exposures in Texas. Isotopic analysis, geochemical data, and paleomagnetic studies indicate that the Cuyania block was adjacent to the Ouachita embayment during Early Cambrian times. The Andean structure shows an east-dipping, basementinvolved thrust overriding the Permian and Tertiary continental deposits. The Cerro Valdivia is a small hill with its lithology

and structure characteristic of the Sierras Pampeanas (Fig. 14). Farther east, Cerro Barbosa, another Grenville-age basement block, has a similar tectonic setting. Both uplifted blocks represent the westernmost exposures of the Sierras Pampeanas basement at these latitudes. Stop 1-4: Eastern Precordillera at Sierra Chica de Zonda The drive is from the San Juan city to the west on Provincial Route 12 in the direction of the town of Zonda, which is some 10 km from the city of San Juan (see Fig. 10). This stop provides the opportunity to examine the Cambrian sequence of limestones bearing the Early to Middle Cambrian trilobites that have been correlated with the Olenellus fauna. The Olenellus fauna of the Eastern Precordillera of San Juan, together with

Evolution of the Pampean flat-slab region

93

68°30′W

SAN JUAN CITY

Legend Tertiary deposits

SIERRA CHICA DE ZONDA

VALLE

30°30′S

(Triangle zone)

DE

ZONDA

Late Paleozoic rocks Early Paleozoic rocks

Co. Barbosa

Precambrian rocks Thrust Active thrust

Neotectonic activity

30°30′S

Rí o

Sa n

Ju

an

STOP 1-3 Co. Valdivia 40

0

5 km

68°30′W

Figure 14. Geologic map of the Eastern Precordillera showing the relationship between the Cambrian and Ordovician deformed limestones of the Sierra Chica de Zonda and the Cerro Valdivia and Cerro Barbosa basement blocks. Note the important neotectonic activity in the foothills of Sierra Chica de Zonda.

the northwestern Scotland and Spitsbergen Island faunas, are the only ones outside of North America. Paleomagnetic, paleoclimatic, isotopic, geochemical, and geochronological data indicate that the limestones and their basement are exotic to Gondwana (see review in Ramos, 2004). The biostratigraphic control and the faunal affinities clearly indicate that an original Laurentian fauna was gradually isolated in Arenigian times, and eventually became endemic by the mid Ordovician, as proposed by Benedetto and Astini (1993) and Benedetto et al. (1999), among others. Gondwanan forms arrived by the Llanvirnian and became dominant in Caradocian to Ashgillian times during the glaciation that dominated this part of Gondwana (see discussion in Benedetto, 2004). Stop 1-4 also allows a regional view of the boundary between the Eastern and Central Precordillera (Fig. 15). This boundary is characterized by a thick-skinned triangle zone (Zapata and Allmendinger, 1996). The thin-skinned, east-vergent Central Precordillera is thrust on synorogenic valley fill of Miocene age. The highly deformed Early Paleozoic rocks of the Eastern Precordillera are uplifted by a basement wedge. These basement structures correlate with a zone of important seismic activity with compressive focal mechanisms at ~14 km depth. The present Juan Pobre valley is an abandoned landscape where the Río San Juan used to flow in Pleistocene times. The latest Quaternary uplift of the Zonda fault displaced the river to the north.

Stop 1-5: Cerro Blanco Farther to the west, after passing 8 km the town of Zonda in the direction to Calingasta is Stop 1-5. The Cerro Blanco dacite has been emplaced synchronously with the uplift and stacking of the Precordillera thrust sequence (Fig. 15). K-Ar dating of these arc-related dacites yielded 6.3 ± 0.7 Ma (whole rock). The abundant basement xenoliths are one of the few indirect evidences of the Grenville-age metamorphic basement directly under the Precordillera. They have been dated by U-Pb in zircons and yielded ages of ca. 1.1 Ga (Kay et al., 1996). The Cerro La Sal dacite exposed in the northern margin of the river, west of Punta Negra (Fig. 15), corresponds to an older volcanic complex with ages around 16 ± 2.9 Ma. The strata of Albarracín Formation (Ta) contain proximal pyroclastic flows that have been produced by explosive collapse of an actively growing dome. Fission-track dating of these tuffs indicates ages from 18 to 7.8 Ma (Vergés et al., 2001). Several intrusive contacts between the sedimentary rocks and the deformed dacites are visible along the road. An eastward pattern of increasingly more depleted signatures is seen toward the east in the Precordillera dacitic magmas. The magmas erupted through crust which, to the east, was less modified by deep crustal thickening and wedge processes. The dacitic melts were related to the Cuyania depleted basement (Kay and Abbruzzi, 1996).

94

Ramos

Figure 15. Geologic map of the lower valley of the Río San Juan showing the triangle zone of Valle de Zonda. Cl and Cu—Lower and Upper Cambrian limestones; Osj—Ordovician San Juan Limestones; Se—Silurian shales; Dt and Dv— Devonian sandstones and shales; Cb—Carboniferous deposits; Ta and Tm—Middle and Late Miocene synorogenic deposits; Tv—arc-related Miocene dacites.

The Tertiary deposits record distal synorogenic facies at the base dominated by shales and fine grained sandstones of Middle Miocene age. Magnetostratigraphic studies recognized an abrupt change in sedimentation rates in the middle to upper section in the Albarracín creek. The strong increase in sedimentation rates at ca. 9 Ma, from 0.09 to 0.44 mm/yr, coincides with a rapid deposition of conglomerates (Vergés et al., 2002). This change is related to the renewed activity in the thrust to the west, and consequent deposition of proximal conglomeratic synorogenic deposits. The distribution and composition of Tertiary magmatism have been related to a shallowing subduction zone and a thickening crust. Isotopic data combined with trace element data from several regions of the shallow subduction zone permit to infer differences in crustal thickening patterns in the development of the modern seismic zone (Kay et al., 1991). Along the Cordillera Principal the thickening and expansion of the magmatic arc is detected prior to 16 Ma, and the end of the subduction-related magmatism occurred at 7 Ma. The Precordillera of San Juan records magmatic activity between 18 and 6 Ma, while in the Sierras Pampeanas volcanism ended at 1.9 Ma (Ramos et al., 1991).

to conglomeratic proximal facies occurs at ca. 4 Ma and indicates an important shifting of the thrust front during Pliocene times. Similar trends have been found in the northern Precordillera by Jordan et al. (1988, 1993, 1997). The Tertiary deposits depict a series of minor normal faults that are related to the lateral ramp that segments the Zonda fault at these latitudes. Stop 1-7: Late Quaternary Lake Deposits Stop 1-7 is 3 km northwest of the previous stop along the road to the town of Ullum (Fig. 16). The lake deposits of the Valentín Formation exposed here are dated at 6500 yr (14C ages) clearly showing that the shore of the Holocene natural lake was even larger than the present lake produced by the dam. A series of neotectonic features are observed along the Zonda fault trace. The Pleistocene terraces have been analyzed by cosmic-ray exposure dating, which has yielded ages of 18.7 and 6.8 10Be ka for the abandonment of the surfaces (Siame et al., 2002). These data indicate a minimum shortening rate of more than 1 mm/yr in the past ~20 k.y., which illustrates the rapid uplift of the Sierra Chica de Zonda (Fig. 16). Displacements along this basement fault may have originated the 1944 San Juan earthquake.

Stop 1-6: Synorogenic Deposits in the Ullum Dam The northern valley of Río San Juan downstream from the Ullum dam exposes Miocene synorogenic deposits along the road to the town of Ullum, 10 km west of the city of San Juan (Fig. 16). A change from distal to proximal facies in these deposits, similar to those previously described, is seen at this stop. The distal facies have been deposited between 8.5 and 4.0 Ma, and the intercalated tuffs have fission track ages of ca. 7 Ma. The change

DAY 2—FROM SAN JUAN TO BARREAL, ARGENTINA Stop 2-1: La Laja Fault Drive north along National Highway 40 heading to the town of Albardón, some 10 km north of San Juan city (Fig. 17). Continue to the La Laja thermal baths heading north

Evolution of the Pampean flat-slab region

95

68°45′W

Tc QTm

Tc Qt

STOP 1-7

40

7 Ma 6,500 k.y.

ULLUM Qt

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n

nd

n Jua

Zo

a Río S

aF

au

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STOP 1-6

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SAN

Figure 16. Geologic map of the active Zonda fault showing the distribution of the lacustrine deposits surrounding the present dam. C-O—Cambro-Ordovician limestones; Tc—Loma de las Tapias Formation (late Miocene); QTm—Mogna Formation (late Pliocene-Quaternary); Qt—Late Pleistocene deposits; Qv— Valentín Formation (Holocene). Slant dashed regions show the locations of the city of San Juan and the town of Ullum; gray square is town of Zonda.

JUAN 12

TcTc

STOP 1-4 31°35′S

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eZ

on

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2 km

for 5.5 km until the first stop. Continue in the direction of the La Laja thermal baths heading north for 5.5 km to reach Stop 2-1 at the La Laja fault. The epicenter site of the large 1944 earthquake which destroyed San Juan city corresponds to a reactivation of La Laja backthrust. This fault places Mio-Pliocene deposits in contact with Pleistocene gravels. The fault has a length of 8 km, a total throw of 18 m, and is part of a system of backthrusts located in the eastern side of the Sierras de Villicum and Rinconada. The fault system extends over 150 km along the eastern border of Precordillera (Bastías, 1990). The 1944 earthquake of magnitude 7.4 produced a reverse displacement of 50 cm on the fault, and the scarp was traceable north of the road affecting recent soils. Most of the thrust system is partially coated by extremely porous travertine carbonates deposited from groundwater activity related to the recent faults. Stop 2-2: Sierra de Villicum Return to the town of Albardón and take National Highway 40 in the direction to Jachal (Fig. 17) for 6 km. to reach Stop 2-2. A major backthrust system is exposed in the southern end of the Sierra de Villicum at this stop. This south-plunging structure has an imbricated backthrust that overrides Cambro-

Ordovician limestones over Miocene strata. The backthrust is the northern segment of the Zonda fault (see Figs. 15, 17, and 18). This succession of Cambrian limestones was the first locality where numerous trilobites of the Early Cambrian Olenellus fauna with North American affinities were found (Borrello, 1963). South of the road there are several abandoned surfaces with ages varying from 6 to 20 ka, dated by Siame et al. (2002). Drive across the Matagusanos Valley Continue north on Route 40 and then head west to the Baños de Talacasto along Provincial Route 436 (see Figs. 10 and 17). The Matagusanos valley is the northern extension of the Zonda triangle zone. It is bounded by the Sierra de Villicum to the east, where a series of west-verging backthrusts of Eastern Precordillera are overriding the synorogenic deposits. Along the valley some active growing structures are seen with west vergence. A wildcat drilled in one of these structures, the Matagusanos anticline, went through more than 6000 m of synorogenic deposits directly overlying the Ordovician San Juan Limestones. The western side is bounded by a series of east-verging imbricated thrusts of the Central Precordillera that are detached in Lower Ordovician limestones, repeating the Early Paleozoic sequence. Figure 18 shows the structure of the triangle zone.

ault da F STOP 2-2

La

La

ja

Fa

ult

Zon

STOP 2-1

Figure 17. Geologic map of Bolsón de Matagusanos showing the west-dipping Zonda thrust fault, the east-dipping back thrust system of the Sierra de Villicum, and the La Laja fault to the east with its recent scarp. COr—Cambro-Ordovician limestones; Tc—Tertiary deposits; Qt— Pleistocene terraces (modified from Furque, 1968).

Eastern Precordillera Sierra de Villicum

Central NW Precordillera

Zonda Faultt Active folding

Sierras Pampeanas

SE

Pie de Palo

Depth (km)

0

10

20

30 Seismicity from Smalley et al. (1993) Precambrian

Seismicity and fault structure after Ramos et al. (2002) CambroOrdovician

Silurian Permian

V=H

Tertiary Quaternary

Figure 18. The triangle zone between Eastern and Central Precordillera with an interpreted structural section of the deep structure of the western Sierras Pampeanas at the Sierra de Pie de Palo (modified from Meig et al., 2006). V = H—vertical and horizontal scales equal.

Evolution of the Pampean flat-slab region Stop 2-3: Sierra de Talacasto Front The thrust front of the Central Precordillera shows three west-dipping forethrusts around the Baños de Talacasto (Fig. 19). Ordovician San Juan Limestones are imbricated over Miocene deposits in the first; Ordovician limestones are over Late Ordovician shales in the second; and Ordovician limestones are over Silurian shales in the third. Accommodation structures such as minor faults and second-order chevron folds are seen in the main thrust sheets. At this stop K-bentonite levels can be observed. These K-bentonites were very significant in the correlation of the Laurentia-derived Cuyania block and Gondwana (Cingolani et al., 1997). Sensitive high-resolution ion microprobe (SHRIMP) ages of zircons in bentonites near this locality yielded ages similar to the magmatic arc rocks of the Famatina region farther east in the Sierras Pampeanas (Fig. 20). The K-bentonites here are older than the typical K-bentonites in the Appalachians. Stop 2-4: Western Sierra de Talacasto Stop 2-4 is a brief stop a short distance northwest of Stop 2-3 to see the thrust to the south that puts Ordovician limestones on Silurian shales (Fig. 19). The stratigraphic separation decreases to the north, where limestones are thrust on similar limestones, indicating out-of-sequence thrusting. Although

97

there is no precise dating on the age of thrusting in this locality, by correlation with adjacent thrusts to the north and south, the timing of deformation is interpreted to be between 10 and 8 Ma (Vergés et al., 2002). The timing of thrusting in the Precordillera shows not only a migration of the thrust front to the east, but also a wave of deformation that youngs to the south. This wave of deformation progressing to the south is also seen in the Sierras Pampeanas uplifts farther east (Ramos et al., 2002), and both are probably linked with the shifting of the Juan Fernández ridge collision along the trench as proposed by Yañez et al. (2002). Drive along the Talacasto Valley and Quebrada de las Burras A few kilometers from Talacasto, leave Route 436 to take a paved road to the southwest toward the town of Calingasta. The road crosses all of the main stratigraphic marine units of the Early Paleozoic of the Central Precordillera (Fig. 21). The 350-m–thick sequence of platform limestones of the San Juan Formation ranges from latest Tremadoc to Early Llanvirnian (Peralta, 2003). An erosional unconformity separates this unit from the clastic deposits of Las Chilcas Formation (uppermost Ordovician– Early Silurian). The Ordovician-Silurian boundary has been recognized south of Baños de Talacasto in Las Chilcas Formation. The yellowish sandstones and shales of the La Chilca Forma-

31°05′S

DE TALACASTO

de Talacast

Quebrada Pob

Quebrada

lete Norte

STOP 2-4

Baños de Talacasto

Legend Quaternary alluvium

o

*

Tertiary deposits Lower Devonian sandstones

STOP 2-3

Silurian shales and sandstones

Que

brad aA

nch a

SIER

RA

* *

*

Ordovician limestomes Thrust

*

Fossil locality

0 68°40′W

500 m

Figure 19. Geologic map of the Quebrada de Talacasto (Talacasto creek) showing the main thrusts of the Central Precordillera at these latitudes (based on Peralta, 2003).

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Figure 20. Comparison of zircon U-Pb ages in Tera-Wasserburg plots of the K-bentonites from Talacasto, central Precordillera on the Cuyania terrane and rhyolitic rocks from the Famatina region along the magmatic arc of Gondwana. Data from Baldo et al. (2003). MSWD—Mean square of weighted deviates.

tion (~110 m thick) of Early to Middle Silurian age unconformably overly the San Juan Limestones in the Quebrada Ancha. The Los Espejos Formation of Late Silurian age is composed of green siltstones and fine sandstones (~250 m thick). Both Silurian units represent the marine foreland distal deposits associated with the collision of the Cuyania terrane with Gondwana. West of the Quebrada Poblete Norte, a complete section of the Talacasto Formation overlies the Silurian deposits. This is the type section of this unit and is represented by olive green sandstones and shales, between 100 and 200 m thick, of Early Devonian age. A thick section of turbidites and shales of the gray color Punta Negra Formation represents a reactivation of the deformation during Middle to Late Devonian times. A deep incised valley of glacial origin cuts the Punta Negra Formation and preserved glacio-lacustrine coarse conglomerates of Late Paleozoic age. Miocene synorogenic deposits unconformably overlie the Early Paleozoic rocks (Fig. 21), a few kilometers west of the junction to Pachaco and Calingasta along the Quebrada de las Burras. This Miocene unit is truncated by the Sasso thrust that repeats Devonian deposits. Stop 2-5: Northern End of Sierra de la Cantera This stop is located in the drainage divide between the Río San Juan drainage basin and the Quebrada de las Burras (commonly dry) river. It is at 2215 m above sea level and corresponds to an active thrust zone related to the La Cantera thrust. This is one of the few active faults in Central Precordillera, where important earthquakes occurred in 1924, 1964, and 1984 (Mingorance, 1998), along many other neotectonic features. This fault which coincides with the Cuesta del Tambolar in the valley of the Río San Juan was interpreted as an out-of-sequence thrust farther south where it truncates overturned Late Paleozoic deposits.

The regional view from this point shows the San Juan limestones in the hanging wall of the Pachaco thrust and the Devonian Punta Negra Formation represented by foreland distal facies. Stop 2-6: The Cerro Blanco de Pachaco After crossing the bridge across the Rio San Juan, turn east toward the small town of Pachaco (Figs. 21 and 22). The Pachaco thrust exposes the westernmost edge of Ordovician platform limestones in the Precordillera. The limestones override Tertiary mudstones and fine sandstones of the Pachaco Formation (Milana et al., 1993). These deposits are distal synorogenic facies probably related to the uplift of the Main Cordillera west of Calingasta at 18–16 Ma (Milana, 1991). The abnormal thickness of the Ordovician San Juan limestones exposed north of the river was explained by Heim (1952) as a splay that repeated the San Juan Limestone (Fig. 22). East of the main peak, with proper sun lighting, a tight syncline which corresponds to the fault trace can be seen. Thick-bedded limestones of the lower member (Late Tremadocian to Early Arenigian) of San Juan Formation are thrust on the thin bedded fossiliferous limestones of the upper member (Late Arenigian to Early Llanvirnian). The competent nature of the limestones contrasts with the highly deformed Silurian shales that are exposed west of Cerro Blanco de Pachaco. Stop 2-7: Ordovician Slope Facies and Olistoliths Continue west along the road paralleling the Rio San Juan in the direction of the town of Calingasta to Stop 2-7 (Fig. 22). The first Ordovician slope facies are represented by the Alcaparrosa Formation, which corresponds to slope and bathyal facies of the Early Paleozoic platform. These rocks host a series of olistoliths up to several hundred meters in

Evolution of the Pampean flat-slab region

99

Figure 21. Drive between Talacasto and Pachaco through the northeast-southwest–trending Quebrada de las Burras. An imbricated fan of thrusts repeats the Devonian sequences (Dv) several times along the route.

diameter of Cambrian and Early Ordovician rocks. At this stop, the olistoliths are mainly limestone blocks of Middle Cambrian age that were deposited by gravitational sliding from the carbonate platform. The size of the olistoliths ranges from a few meters up to a couple of kilometers. There are many of them along the edge of the Early Paleozoic platform in this region.

ages of the Granite-Rhyolite province west of the Grenville Front in the northwestern corner of the Ouachita embayment (Thomas et al., 2000). Mafic rocks are interbedded with this sedimentary sequence on the western side of the Sierra de Tontal. As a whole, this sequence has been interpreted as a klippe of western facies thrust on to the platform edge. The Andean age west-dipping Los Ratones thrust fault at this stop (Fig. 22) is covered by Quaternary deposits.

Stop 2-8: Sierra de Tontal Drive from Western Precordillera to Barreal Continue west along the road paralleling the Rio San Juan in the direction of the town of Calingasta to Stop 2-8 (Fig. 22). Most of the Sierra de Tontal is composed of the Alcaparrosa Formation. These slope facies vary from coarse proximal turbidites to more distal ones. The graywackes and shales bearing scarce graptolites are tightly folded and developed an incipient cleavage that increases to the west. The dense debris flows with resedimented conglomeratic blocks seen at this stop were derived from the synrift facies. Some pebbles in equivalent deposits farther to the north have been dated at 1367 ± 5 and 1370 ± 2 Ma (U-Pb in zircons). These ages are in concordance with the 1.3–1.4 Ga

The road continues to the west along the southern margin of Rio San Juan heading to the town of Calingasta ~38 km farther west. Turn south at Calingasta to spend the night in Barreal. The western section of the San Juan valley is characterized by a complex series of thrusts, which repeat the different Early Paleozoic slope deposits that are tectonically interfingered with more oceanic facies. The valley widens at Calingasta, in the junction of Los Patos and Castaño rivers. There, the Uspallata-Calingasta valley, a north-trending depression, probably corresponds to an unbroken piggy-back basin, similar to the one described farther

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Figure 22. Map of the classical section of the Cerro Blanco de Pachaco at Stop 2-6 (after Heim, 1952) and the region of the subsequent Stops 2-7 to 3-1 in the Western Precordillera.

north in this depression by Beer et al. (1990). Along the road isolated patches of Triassic synrift deposits are observed. These deposits are controlled by northeast-trending half grabens developed in the hanging wall of the suture between Cuyania and Chilenia, which at these latitudes coincides with the boundary between Precordillera and Cordillera Frontal. This depression has many similarities with the Rocky Mountain trench of North America. DAY 3—FROM BARREAL TO USPALLATA, ARGENTINA The Barreal area has many classic localities studied by Du Toit, Keidel, and others, since the beginning of the last century. In the area, there are excellent exposures recording the Late Paleozoic Gondwanian glaciation with glaciated pavements, dropstones, striated clasts, till, and outwash deposits. These deposits were used as a confirmation of the importance and extension of the Late Paleozoic ice cap (Keidel, 1916; Du Toit, 1927). The evidence of this Late Paleozoic glaciation described by Keidel (1916) was one of the first-order pieces of geological evidence used by Wegener (1929) to hypothesize the continental drift. The drive from Barreal to Calingasta provides an opportunity to observe the magnificent scenery of snowy mountains along the eastern margin of the Cordillera Frontal. If weather conditions permit, the northern side of the Mount Aconcagua far to the south can be seen from the road.

tance of ~5 km along the road, two west-verging thrusts (Tontal thrusts) on the eastern side of an Andean age triangle zone and two east-vergent thrusts (western Agua de los Pajaritos thrust and eastern Carrizal thrust) on the western side can be seen. A short walk east of the Carrizal thrust and west of the IMSA building provides access to an important angular unconformity between heavily deformed Devonian flysch and less deformed Carboniferous foreland deposits. These Carboniferous continental deposits bear an abundant Early Carboniferous Rhacopteris flora typical of the Gondwanian realm. They also have rhyolitic and granitic pebbles derived from the Chilenia terrane. This unconformity is related to the docking of the Chilenia terrane against the protomargin of Gondwana during the Late Devonian times. Stop 3-2: Low-Grade Metamorphic Rocks Follow the road west along the Rio San Juan to Stop 3-2 on Figure 22. The low-grade metamorphic rocks at this stop belong to the Don Polo Formation, a unit that has been traditionally considered as the oldest exposed in the area. As it is in tectonic contact with other Ordovician rocks, its age is in dispute between Late Proterozoic (?) and Ordovician. The sequence consists of highly deformed graywackes and shales. To the south, the unit grades to anchimetamorphic to low-grade metamorphic facies. This assemblage has been correlated with the Taconian allochthon of the Appalachians by Nullo and Stephens (1996).

Stop 3-1: Km 114 Triangle Zone

Stop 3-3: Ordovician ophiolites

This stop is in the region of Km 114 on the road west from Calingasta in the direction of Pachaco. The stop and the important features at this stop are shown on the map in Figure 22. In a dis-

Continue west along the Rio San Juan to Stop 3-3 on Figure 22. Beautiful basaltic pillow lavas crop out on the south side of the road near where the road bends and the high cordilleras are

Evolution of the Pampean flat-slab region

Stop 3-4: El Alcázar This stop is located south of the city of Calingasta, ~17 km north of Barreal (Fig. 23) near the locality of Hilario. The small outcrop of Triassic sag phase deposits at this stop is characterized by tuffs, reworked tuffs, siltstones, and fine-grained sandstones of lacustrine and fluvial facies. These sediments were deposited along the western and passive ramp of a half-graben system (López Gamundi and Astini, 1992). The active margin is located to the east and is characterized by coarse conglomerates and red sandstones similar to the synrift deposits of the Río Mendoza Formation near Potrerillos dam (see chapter 3, this volume). Although the age of the deposits in this locality is poorly constrained, other outcrops along the Río de los Patos valley have abundant flora and palynomorphs indicating a Late Triassic age. Drive along the Calingasta-Uspallata Valley The long depression that separates the Western Precordillera from the Frontal Cordillera coincides with an old tectonic feature—the Early Paleozoic western continental margin of first Cuyania and then Gondwana. This depression is known as the Iglesia-Calingasta-Uspallata valley and extends for over 300 km (see Fig. 23). This drive from Calingasta to the south for 147 km to the town of Uspallata shows the clear differences between the Precordillera and the foothills of the Cordillera Frontal. On the western side, the black tones of the Carboniferous deposits contrast with the light pink colors of the Gondwanian granitoids. The structure is homogeneously massive in comparison with the different thrust sheets observed in the Precordillera. Along the eastern side of the road, slope facies of Ordovician and Silurian age are interbedded with basaltic pillow lavas. Alum

CALINGASTA USPALLATA VALLEY

Late Paleozoic granitoids

30°S

Rodeo

C O R D I L LE R A F R O N T A L

first seen to the west. The basalts are at the base of the Alcaparrosa Formation, and upward the slope facies are interfingered with them. Locally, they are associated with fine-grained cherts. This stop shows some of the best preserved Ordovician pillow lavas in the Precordillera. A detailed observation of the pillows shows the smectitic “bread” crust, amygdules, and radial cleavage. Although the outcrop here only shows pillow lavas and breccias, farther to the east and to the south there are outcrops of gabbros, peridotites, and sills typical of an ophiolitic assemblage. This assemblage is exposed along a belt of more than 1000 km that represents the suture between the Cuyania and Chilenia terranes. Secondary mineralization is due to both sea water alteration and later tectonic metamorphism. Their geochemistry indicates the strong oceanic affinities of an enriched mid-oceanic ridge basalt (E-MORB) setting, which is similar to other oceanic tholeiites of abnormal ridges. Geochemical and isotopic characteristics are similar to the Newfoundland ophiolites (Kay et al., 1984; Haller and Ramos, 1984). This stop provides a cross section of the pillows. If time allows, an optional stop on the west of the bridge at Calingasta Bridge (Fig. 22) provides a plane view of the ophiolites.

101

Pismanta Las Flores

Puchuzúm

32°S

Calingasta STOP 3-4

SAN JUAN

Barreal

STOP 3-5 STOP 4-1,2

STOP 4-3

70°W

Uspallata 0

50 km

MENDOZA

Figure 23. The Calingasta-Uspallata valley with location of Stops 3-4 and 3-5 in the valley, 4-1 in the “Darwin forest,” and Stop 4-2 at Cerro Colorado. See Figure 24 for further geologic details.

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mines exploit secondary minerals associated with the oceanic facies. Along the eastern side of the road, outstanding white outcrops of Triassic rocks are composed of siltstones and tuffs containing abundant flora remains of Late Triassic age. The sequence comprises thick red conglomerates and sandstones of the synrift facies at the base, and sag deposits with siltstones, shales, and tuffs in the upper section. All of these small basins are developed on the hanging wall of the suture between the Cuyania and Chilenia terranes. South of the flat basin of the Barreal de Leoncito south of Barreal, the valley is deflected to the east with an echelon pattern, and shows a more complex broken structure. Several isolated exposures of Early Paleozoic rocks are seen in the valley axis. Some of these rocks are imbricated SiluroDevonian strata and mafic rocks with west vergence in the Sierra de Sandalio (Cortés, 1989). The Sierra de Sandalio is located west of Sierra de Cortaderas, where the main outcrops of ophiolites are exposed (east of Stop 4-1 on Fig. 23). On the western side of the valley, thick sequences of Choiyoi deposits are interbedded with sedimentary lacustrine facies (Cortés, 1985). Several half-graben systems have been identified along this border of Cordillera Frontal. Important Permian magmatism is responsible for the porphyry copper system present near Yalguaraz and known as the San Jorge prospect. Upon arriving in the town of Uspallata, the surface of the uplifted peneplain of the Cordillera del Tigre is seen at 4500–5000 m high all along the western margin of the valley. Stop 3-5: Agua Hedionda Stop 3-5 is a little more than half way to Uspallata from Calingasta (Fig. 23); the latitude and longitude are 32°01′15″S and 69°21′10″W. Highly deformed Devonian rocks of the Ciénaga del Medio Group are exposed at this stop. These graywackes and shales represent distal turbiditic sequences characteristic of the slope facies that are interbedded with pillow lavas and mafic dikes farther south. The geochemistry and isotopic signature of the magmatic rocks indicate an oceanic source. Plant remains point to a possible Silurian or Devonian age for these deposits (Cortés, 1993; Cortés and Kay, 1994). These rocks are either part of the accretionary prism of the eastern edge of the Chilenia terrane or part of the subduction complex of the Cuyania terrane. Optional Cultural Stop: Tambillos In the foothills of Cordillera Frontal, isolated segments of the Camino del Inca (Inca trail) can be observed as in Tambillos. In this locality just north of Uspallata, there are preserved ruins of a small setting built by the Incas in the fifteenth century, to be used as lodging for the messengers that crossed up and down the empire. There are several of these places along the Río Mendoza valley set apart at ~50 or 60 km distance. Picheuta and Ranchillos are among them. After this stop, continue south to the town of Uspallata.

DAY 4—USPALLATA REGION AND WEST TO PUENTE DEL INCA The first two stops on Day 4 are to see the Miocene volcanic rocks and Triassic basalts and sedimentary sequences east of the town of Uspallata. Stop 4-1: Darwin Forest From the town of Uspallata, proceed north on old Route 7 in the direction of Villavicencio and the city of Mendoza. Stop 4-1 is a few kilometers east of the Agua de la Zorra (Fig. 24) and can be located on the south side of the road fragments of a broken sign. This is the classic locality where Charles Darwin found the first fossil floras of southern South America in 1835. He reported a fossil forest with vertical silicified stems of “Araucarites” in life position near Agua de la Zorra (Fig. 24). Electron microscope studies performed in recent years confirm that the trunks belong to Araucarioxylon (Brea, 1997). The fossil horizon is in the Potrerillos Formation of Late Triassic age. The red sandstones and siltstones of the Potrerillos Formation are interfingered with within-plate alkaline basalts of middle Triassic age (235 Ma, Ramos and Kay, 1991). Columnar jointed basaltic sills are well seen along the road near Agua de la Zorra (see horizontal striped pattern on map in Fig. 24) and can be observed on the way to Stop 4-2. Stop 4-2: Cerro Colorado Dacite Continue east on old Route 7 (Fig. 24). The volcanic center of Cerro Colorado is part of a series of dacitic and andesitic bodies of early Miocene age, which are widespread along the Precordillera of San Juan and Mendoza. They erupted prior to the shallowing of the Benioff zone in Miocene times at these latitudes (Fig. 24). Cerro Colorado has been dated by K-Ar whole-rock analyses yielding an age of 18.4 ± 0.7 Ma (Kay et al., 1991; Ramos et al., 1991). Geochemical and isotopic studies indicate a mixed mantle and crustal source with low La/Yb ratios indicating magma genesis prior to the thickening of the crust at these latitudes (Kay et al., 1991). Drive West from the Town of Uspallata into the Frontal and Main Cordilleras The drive to the west of the town of Uspallata town along National Highway 7 for 80 km provides the opportunity to cross the Cordillera Frontal along the northern and southern margins of the Rio Mendoza Valley. The drive features the relationship between Late Paleozoic subduction and the deformation known as the Gondwanide orogeny since the early works of Keidel and Du Toit, as well as the subsequent generalized extension responsible of the Choiyoi Permo-Triassic magmatic province (Kay et al., 1989). These features are clearly seen across this sector of the Andes along the boundary between Precordillera and Cordillera Frontal. The Cordillera Frontal is the locus of Late Paleozoic tectonics

Evolution of the Pampean flat-slab region

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69°15′W

PARAMILLOS DE USPALLATA 0

Cerro Colorado

3 km

Miocene volcanics

Uspallata Group STOP 4-2

Mina de Los Paramillos Agua de La Zorra 32°25′S

STOP 4-1

Basalts

32°25′S

Cruz de Los Paramillos

Figure 24. Geologic map of a sector of the Cuyo rift, with location of the Triassic “Darwin Forest” at Stop 4-1 and Miocene volcanic rocks at Cerro Colorado at Stop 4-2. The horizontal line patterns indicate the localities of Triassic basaltic rocks.

and is where the Choiyoi province exposes the subduction related volcanics, the San Rafael orogenic phase, widespread rhyolitic volcanism, and coeval Triassic Cuyo rift deposits. The Cordillera Frontal uplift was one of the first consequences of the collision of the Juan Fernández ridge at this latitude and the development of the present Pampean flat slab. The late Miocene peneplain is still preserved in the Cordillera Frontal. The following stops supplement those along this route in chapter 3 of this volume.

are recognized by their homogeneous texture and their columnar jointing. At this stop section, most of the ignimbritic and pyroclastic flows are of Triassic age. Similar rhyolitic domes have been dated at 203–205 Ma farther west (Latest Triassic). Stop 4-5: Polvaredas Normal Fault

At this stop along Route 7 (31°36′53.5″; 69°23′54.1″; 1874 m asl), the southern end of the Calingasta-Uspallata valley can be observed (Fig. 25). The Cordillera Frontal at these latitudes is thrust over the Precordillera. A series of imbricate thrusts repeats Choiyoi volcanic rocks on top of Miocene continental foreland basin strata. Based on the magnetostratigraphic study performed in the synorogenic deposits related to the Cordillera Frontal uplift, these thrust initiated at ca. 9 Ma (Irigoyen et al., 2002). This is immediately after the collision of the Juan Fernández ridge against the trench.

This stop provides a good example of the Gondwanian tectonics (location on map in Fig. 25). A normal fault puts the western facies of the Carboniferous deposits in contact with the pyroclastic sequences of the Choiyoi Group. The fault trace was intruded by a rhyolitic dike of Late Triassic age, which excludes an Andean reactivation of the fault. The well-stratified pyroclastic deposits of the Choiyoi Group have been dated here at 235–238 Ma by K-Ar (Pérez and Ramos, 1996), which indicates a Middle Triassic age. Note that the rhyolitic volcanism of this region is synchronous with the alkaline basalts at the Agua de la Zorra (235 Ma). The extensional regime that controlled the Choiyoi rhyolites was coeval with the half graben developed farther to the east in the modern foreland area.

Stop 4-4: Picheuta Rhyolitic Porphyry

Stop 4-6: San Rafael Orogenic Phase

This stop is in the middle of the Choiyoi province at these latitudes (Fig. 25). Thick piles, up to 2–4 km thick, of pyroclastic and volcanic rocks of rhyolitic composition represent a widespread extensional period. Several rhyolitic domes north of Río Picheuta are emplaced in the pyroclastic sequences. These domes

The angular unconformity that separates the Late Carboniferous and Early Permian turbidites from the Choiyoi Group represents the San Rafael diastrophic phase of Middle Permian age (location of stop on map in Fig. 25). The sedimentary rocks have contact metamorphism produced by granitoids of Early

Stop 4-3: Southern End of the Uspallata Valley

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Figure 25. Generalized map of the High Andes from Uspallata to the Chilean border indicating stop locations. Stippled pattern—Late Paleozoic strata; crosses—Early Permian granitoids; Trch—Choiyoi volcanics; Tv—Miocene volcanic rocks; Tsm—Santa María conglomerates.

Permian age (ca. 287 Ma). The Cordillera Frontal between Punta de Vacas and the Uspallata valley has a simple Andean structure, characterized by a large uplifted block of basement. All the Late Paleozoic–Triassic rocks of the Cordón del Plata, which is the southern expression of the Cordillera Frontal (see section in Fig. 26), have been passively transported to the east in a detachment located on the Carboniferous shales. Stop 4-7: Río Colorado Granite This stop shows the roof of the Lower Permian granite intruded in the Carboniferous sequence (Fig. 25). The granite was first described by González Bonorino (1950b) to illustrate the stopping mechanism in different stag in the emplacement of blocks of Carboniferous rocks. There are blocks partially detached falling into the granite and blocks partially absorbed.

There are several other granitic bodies between this stop and Punta de Vacas, all of which are preserved at shallow levels with their roofs partially exposed. These granitoids are subductionrelated and were part of the Early Permian magmatic arc. Stop 4-8: Volcán Tupungato At this stop (Fig. 25), the Tupungato volcano (6800 m) can be seen looking to the south along the Rio Tupungato. This is the first Quaternary volcano south of the flat-slab segment of the central Andes, and is located ~40 km south of Punta de Vacas. The last activity recorded in this volcano is a pyroclastic flow in its northern slope of 0.7 ± 0.3 Ma. The Tupungatito volcano located nearby the main center is active and has fumaroles. The linear trend of Río Tupungato is also controlled by a Gondwanian fault.

E s

t aul

La Fa s Ca ult rre ra

CHILE

ARGENTINA

nite

nte

sF

CORDILLERA FRONTAL

Pe

aul t

CORDILLERA PRINCIPAL

ro F Po cu

C DE ORD LA ILL CO ER ST A A

W

0

Figure 26. Schematic cross section of the Andes between the Uspallata Valley and the Cordillera de la Costa (32°30′S).

10 km

Evolution of the Pampean flat-slab region Optional Cultural Stop: Punta de Vacas Shelter This is one of the four shelters on the Argentine side that are preserved from the late eighteenth century. They were built by the Correos Reales de España (Spanish Royal Mail) in 1765. In this shelter, Charles Darwin overnighted during his crossing to Paso de la Cumbre in 1835 (see Fig. 2, and also Darwin [1846] for a picturesque description of the place). Stop 4-9: Cruz de Caña Granite This stop shows a subduction-related granitoid of Early Permian age, emplaced in Late Carboniferous black shales and graywackes, and truncated by extensional faults (Fig. 25). These normal faults do not affect the Middle Jurassic limestones that unconformably overlie the Carboniferous rocks. DAY 5—FROM PUENTE DEL INCA, ARGENTINA, TO VIÑA DEL MAR, CHILE One of the outstanding features of the Mesozoic succession is the development of many Mesozoic marine sequences controlled by Pacific transgressions and regressions. Those sequences are grouped into four sedimentary cycles which are separated by regional first-order unconformities. These sequences are clearly depicted in the Neuquén Basin farther south where a Liassic to Neocomian marine retroarc to foreland basin developed is well developed behind the magmatic arc in the eastern foothills of the Cordillera. The Neuquén basin is linked to the north with the Aconcagua basin, which has a different paleogeography with a larger participation of volcanic rocks, but similar stratigraphic cycles in the middle Jurassic to Cretaceous evolution. There are interesting parallels between the evolution of the basin and the magmatic arc history. The main periods of regional unconformities in the foreland are coincident with the times at which the magmatic arc migrated eastward. The intermittent nature of the magmatic activity as well as the spatial variation of the volcanic front in the Andes appears to be closely related to

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changes in plate motion controlled by variation in the spreading velocities of the Pacific oceanic ridges. Forearc subduction erosion cannot explain by itself the magnitude of the shifting of the Neogene volcanic front. Most of the Chilean and the westernmost Argentine Mesozoic basins of the Cordillera Principal are intra-arc basins controlled by the development of two distinctive arcs: an inner, more active arc along the Cordillera de la Costa in Chile, where the main andesitic activity was concentrated, and an outer arc that produced rock suites mainly of andesitic to bimodal composition. Several authors have proposed that an extensional regime in the arc massif region mainly during the Early Cretaceous was responsible for the intra-arc basin development (Fig. 27). This process has been explained by the negative rollback velocity of the South American plate previous to the opening of the South Atlantic Ocean. The intra-arc basins were active until Early Aptian times when an important eastward migration of the main magmatic arc occurred, together with a sea-level lowstand in the retroarc and intra-arc basins and the development of a single and expanded central arc. The retroarc easternmost basin was exclusively continental from this time on, and the Pacific seas no longer reached the eastern side of the cordillera. This important paleogeographic change is closely linked with the beginning of one of the periods of higher spreading rates within the South Atlantic and the change to a positive rollback velocity of the South American plate. The maximum sea-floor spreading rate at the Pacific and the South Atlantic spreading centers was reached during the Late Cretaceous (ca. 80–110 Ma) and may be responsible for a new stage in the Andean evolution. At this stage (Fig. 27), a mountain chain was built and the deformation of the strata on the eastern flank of the orogen formed a fold-and-thrust belt. As a consequence, a foreland basin developed at the leading edge of the deformation, due to tectonic loading of the adjacent thrust belt. At the final stage of compression, several granitoid stocks were emplaced in the arc massif.

Figure 27. Lithospheric scale cross-section model for intra-arc extension developed in the Cordillera Principal in the Early Cretaceous (after Ramos, 1989).

Figure 28. Cross sections of the Aconcagua fold-and-thrust belt near Puente del Inca looking to the north (after Ramos, 1985b). West is to the left.

Evolution of the Pampean flat-slab region As a concluding remark on the orogenic cycle, it is necessary to emphasize that the different paleogeographic settings, the major depositional sequences and the successive tectonic regimes seem to be controlled by the subduction dynamics as inferred by the changes in the tectonic regime and from the migration of the volcanic front. The latest change in the subduction geometry was controlled by the docking of the Juan Fernández ridge (Pilger, 1984, and subsequent works).

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All these sequences are intruded by thick Miocene dikes of trachyte (15 Ma). In the upper thrust plates, basaltic and andesitic lenses can be seen interbedded into the sequence. These volcanic rocks are more abundant in the westernmost thrust sheets. Stop 5-3: Rio Horcones and Cerro Aconcagua

This is the thrust front of the Aconcagua fold-and-thrust belt (Figs. 28–31). Middle to Late Jurassic limestones override the thick Santa María conglomerates of early to middle Miocene age. These conglomerates unconformably overlie marine and continental Jurassic strata. The Penitentes thrust dips from 50° to 22° to the west. Erosion is almost forming a klippe on the top of Cerro Visera. Figure 29 illustrates the 22° west-dipping footwall ramp that this thrust developed in the Santa María Conglomerates in the southwestern slope of Cerro Penitentes (not seen from the road).

Continue west along National Highway 7 heading west until the Parque Provincial Aconcagua. At this stop (Figs. 30 and 32), there is a magnificent view of the Pared Sur (south wall) of Cerro Aconcagua (6967 m asl). The wall is formed by volcanic and breccia flows of andesitic composition of the Aconcagua Volcanic Complex (15–9 Ma, Ramos and Yrigoyen, 1987; Godoy et al., 1988). On the western side of the valley, there is an imbrication of Jurassic continental red beds and Early Cretaceous limestones. On the eastern side, the diapiric effects of the Late Jurassic gypsum of the Auquilco Formation produced the complex structure of Cerro Panta. To the south, the imbrication of the second to fourth thrust sheet repeating the different Titho-Neocomian units can be observed.

Stop 5-2: Puente del Inca

Stop 5-4: Quebrada Navarro

This is the most classic section of the High Andes. Compare the evolution of thought since Darwin and Schiller by comparing their sections (Figs. 2 and 3) with the present interpretation (Figs. 28 and 31). Figure 31 shows the autochthon represented by Carboniferous hornfels, thin pyroclastic deposits of Choiyoi Group, and Jurassic limestones and conglomerates of proximal facies, and small outcrops of marine Early Cretaceous deposits. The first thrust sheet is composed of Middle to Late Jurassic marine and continental deposits and Early Cretaceous continental to transitional marine deposits. The second thrust sheet is represented by Late Jurassic gypsum at the base and TithoNeocomian continental and marine strata.

At this altitude, the Quebrada de Navarro thrust, which dips more than 80° to the west can be seen (Fig. 32). This is the result of the rotation of the different thrust sheets. It was previously interpreted as an original high angle fault. Based on that interpretation, Zeil (1979) characterized the Andes as being formed by typical upthrusts (high-angle thrusts). A detailed examination of the cut-off angles in the different thrusts points out to lower angles (~20°–22° at maximum).

Stop 5-1: Cerro Penitentes Thrust Front

Optional Cultural Stop: Paramillos de las Cuevas Shelter An optional stop can be made to visit a colonial shelter built in 1765. The Spanish Royal Mail constructed those shelters to make it possible to cross the Main Andes during the winter, because at that time it was very unsafe to cross the Magellan strait due to the war with the British. The Paramillo de las Cuevas shelter (“casucha”) was built by the Spaniards as one of eight shelters across the Andes. It has a prolific history with many outstanding visitors in the nineteenth century, such as D.F. Sarmiento and the San Martin army during the cross of the Andes during the independence times. Several famous travelers described the penuries related to the winter and the snow storms and strong winds (see Darwin stories in his famous trip across the Andes in 1835). Stop 5-5: Las Cuevas Rock Avalanche

Figure 29. Map of the thrust front of Cerro Penitentes looking to the south. West is to the right. Abbreviations as in Figure 30.

The Neocomian limestones in the thrust seen here override Late Jurassic–Early Cretaceous red beds. The thickness of the continental and volcaniclastic deposits exceeds several times the normal thickness of these units, indicating their proximity to the volcanic arc. The limestones bear Olcostephanus sp. of Late

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Figure 30. Geologic map of the Aconcagua fold-and-thrust belt. Cat—Carboniferous Alto Tupungato Formation; Trch—Volcanics of the Choiyoi Group; c—Early Jurassic breccia; Jlm—Jurassic La Manga Formation; e—Gypsum of Auquilco Formation; Jt—Red beds of Jurassic Tordillo Formation; JKvm—Tithono-Berriasian Vaca Muerta Formation; Km—Sandstones of Valanginian Mulichinco Formation; Ka—Limestones of Hauterivian Agrio Formation; Kd—Red beds of Barremian Diamante Formation; Kv—Early Cretaceous volcanics; l—Miocene dacites of Puente del Inca; m—Early Cretaceous diorites; Tsm—Miocene Santa María Conglomerates; Qgl and Qsl—glacial and slump deposits; Q— undifferentiated alluvial sediments.

Valanginian age (Aguirre-Urreta and Rawson, 1997). A spectacular rock avalanche was produced from one of the volcaniclastic members of the Mesozoic deposits. Stop 5-6: Matienzo Valley (Argentina) A complex out-of-sequence thrust folded and thrust the red beds and volcaniclastic rocks of the Early Cretaceous Cristo Redentor and Juncal Formations over the Neocomian limestones. The thick pile of Early Cretaceous rocks is deformed in an anticline structure. To the north, along the Cordillera del Límite, interbedded limestones and volcanics are representing the Neocomian deposits in the westernmost thrust sheet.

The Cordillera Principal at the Chilean Slope The drive after crossing the international tunnel to Chile along the highway in the direction of the city of Viña del Mar will show the western slope of the Andes. The Cordillera Principal corresponds to the highest Andean peaks along the ChileArgentina border. The Chilean side (see Figs. 26 and 33) is made up of gently folded Mesozoic (Early Cretaceous) and Tertiary (Paleogene and Neogene) volcano sedimentary formations intruded by Tertiary (Miocene) granodioritic stocks. The whole sequence is thrust eastward over the Jurassic–Early Cretaceous sedimentary formations of the Aconcagua fold-and-thrust belt between Las Cuevas and Puente del Inca. The Pocuro fault

Evolution of the Pampean flat-slab region

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Figure 31. Cross section of the panoramic view of the Puente del Inca section looking to the south with the modern structural interpretation. Crt—Alto Tupungato Formation (Carboniferous); Trch—Choiyoi Group (Triassic); Jm—La Manga Formation (Middle Jurassic); Jt—Tordillo Formation (Late Jurassic); JKi—Early Titho-Neocomian deposits; JKm—Middle Neocomian deposits; JKs—Late Neocomian deposits; Tc— Santa María conglomerates (Miocene); Tv—Puente del Inca Trachyte (Miocene).

STOP 55-5

STOP 55-3 STOP 55-4

STOP 55-2

STOP 55-6 STOP 5-1

Figure 32. Structural sketch map of the Aconcagua fold-and-thrust belt with locations of Stops 5-2 to 5-6.

(Fig. 26) is a major structural discontinuity that separates the Cordillera Principal from the Cordillera de la Costa. The main components of the geological history of the Chilean part (Fig. 33) of this segment of the Andes (Segment B from 27° to 33°S of Mpodozis and Ramos, 1990) are as follows: (1) emplacement along the coast of the Pichidangui suspect ter-

rane; (2) development in the Jurassic of a subduction related magmatic arc (Cordillera de la Costa) and of a retroarc sedimentary sequence toward the foreland (Aconcagua platform); (3) during the Early Cretaceous, a large subsiding volcanic zone, the Central Chile “aborted” marginal basin or intra-arc basin that formed behind the inner arc; (4) this basin tectonically collapsed during

STOP 5 5-9 STOP 5 5-8

STOP 5 5-7

Figure 33. Geologic map and cross section of the Aconcagua River area modified from Rivano et al. (1990). Symbols are: (1) Quaternary alluvial deposits; (2) MiocenePliocene coastal marine sediments; (3) Miocene Farellones volcanics; (4) Late Cretaceous Vinitas Formation; (5) Late Cretaceous Pelambres Complex (Abanico and Juncal Formations); (6) Cretaceous Mendoza Group; (7) Cretaceous Lo Valle Formation; (8) Cretaceous Las Chilcas Formation; (9) Lower Cretaceous Veta Negra Formation; (10) Lower Lo Prado Formation; (11) Jurassic Cerro La Calera Formation; (12) Jurassic Aijal Formation; (13) Triassic–Jurassic El Cajon Formation; (14) Paleozoic metamorphic rocks of the Quintay Formation; (15) Upper Miocene granitoids; (16) Oligocene-Miocene; (17) Cretaceous; (18) Jurassic; (19) Paleozoic. The schematic cross section between Llaillay and Portillo after Rivano et al. (1985) illustrates the increase of deformation from west to east. Symbols are: (1) gypsum; (2) carbonate rocks; (3) conglomerates and sandstones; (4) volcaniclastic rocks; (5) andesites and tuffs; (6) rhyolites; (7) intrusive granitoids; (8) cataclastic rocks.

STOP 5-13

STOP 5-12

STOP 5-11

STOP 5-10

Evolution of the Pampean flat-slab region the Late Cretaceous; (5) in the Cretaceous-Tertiary, the area witnessed a progressively eastward migration of magmatic activity and deformation; and (6) magmatism vanished in the Late Miocene as a consequence of the shallowing of the subduction angle. The Pichidangui terrane of unknown size seems to have been emplaced by margin-parallel, strike-slip motion in the Late Triassic–Early Jurassic (Forsythe et al., 1986). Strike-slip motion along the Chilean coast was accompanied in Argentina, by the development of Triassic rift basins (Ramos and Kay, 1991). They succeeded Choiyoi magmatism at the end of the Gondwanide history of this Andean segment. The oldest in situ post-Pichidangui sequences in the Coast Cordillera are the Ajial and Cerro Calera Formations that consist chiefly of rhyolitic and pyroclastic rocks. They erupted during an interval of Early Jurassic volcanism (Piracés, 1977). Volcanic eruptions were partly submarine as indicated by the presence of interbedded Bajocian limestones (Fig. 32). At the same time, large gabbroic to granitic plutons were emplaced in the coastal region (Rivano et al., 1985). These plutons have yielded K/Ar ages ranging from 191 to 138 Ma (Munizaga and Vicente, 1982; Rivano et al., 1985). In the retroarc region to the east, a Liassic to Dogger marine transgression initiated as a clastic-carbonate platform that developed over the Late Paleozoic igneous basement of the Cordillera Frontal. Uplift during the Late Jurassic caused produced subaerial conditions, which are indicated by the Horqueta Formation in the Cordillera de la Costa (Piracés, 1977). In the retroarc region, an Oxfordian marine regression is suggested by thick evaporite units such as the Auquilco gypsum. These units are overlain by upper Jurassic continental red beds that were derived from erosion of the volcanic rocks of the Cordillera de la Costa. The sea advanced over the Aconcagua platform during the Early Cretaceous (Ramos, 1985a, 1985b). This Early Cretaceous event was followed by: (a) a noticeable decrease of plutonism in the Cordillera de la Costa, where only small stocks are found just to the east of the Jurassic magmatic belt (Nasi, 1984; Rivano and Sepúlveda, 1986); (b) a progressive volcanic change from andesites to basalts, interbedded with marine carbonate rocks (Lo Prado and Veta Negra Formations) having low 87Sr/86Sr initial ratios (Piracés, 1976; Levi and Aguirre, 1981; Aberg et al., 1984); and (c) eastward expansion of volcanism into the Aconcagua platform in which the Early Cretaceous Los Pelambres Group lavas interfinger with Neocomian limestones along the ChileanArgentine border. Aberg et al. (1984) and Levi and Aguirre (1981) interpreted the widespread Early Cretaceous volcanic activity as taking place in a short-lived “aborted” marginal basin of Hauterivian-Albian age. The term aborted was used by Aberg et al. (1984) to denote a zone where large volumes of basalts and andesites were erupted through a thin, attenuated continental crust which did not evolve into an oceanic crust floored basin. In central Chile, a high thermal gradient associated with subsidence caused pervasive “burial” metamorphism of the volcanic pile. At the end of the Early Cretaceous, the basin was filled with

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coarse red conglomerates interbedded with lavas and limestones (Las Chilcas Formation, Rivano et al., 1985). Sometime during the Late Cretaceous, the basin collapsed and subsequent deformation led to the beginning of the thrusting of the Cretaceous volcanic pile toward the east over the sediments of the Aconcagua platform. Large Late Cretaceous batholiths were emplaced along the axis of the volcano-tectonic rift. Despite the lack of a clear record of Late Cretaceous sediments or volcanism, a belt of Paleocene granitoids was emplaced along the western foothills of the Cordillera Principal. An even younger belt (Miocene) of granodioritic stocks intruded farther to the east, linked to the Miocene volcanics of the Farellones Formation (Mpodozis and Ramos, 1990). This unit, which overlies the Early Cretaceous volcanics of the Pelambres Group, represents the youngest volcanism recorded in the Cordillera Principal. Miocene shallowing of the subduction angle resulted in the shut-off of the magmatism (Jordan et al., 1983a, 1983b; Kay et al., 1987) and beginning of deformation in the Precordillera and Sierras Pampeanas. The route in Chile (Figs. 26 and 33) goes down the Aconcagua valley crossing the Cordillera Principal, the Pocuro fault, and the Cordillera de la Costa (Portillo to Los Andes to San Felipe), and then heads to La Calera and ends at Concón and Viña del Mar along the Late Paleozoic to Jurassic coastal batholith in the Papudo-Valparaíso area. The stops locations are shown on Figure 33. Stop 5-7: Laguna del Inca View to the north of the lake (Laguna del Inca) and the steeply dipping strata of the Juncal Formation (infill of the Early Cretaceous central Chile intra-arc basin) along the international boundary between Chile and Argentina. On the eastern side, Neocomian limestones are interbedded with volcaniclastic and volcanic sequences of the Pelambres Complex (Morata and Aguirre, 2003). Farther to the west, a Miocene diorite intrudes the Juncal Formation. To the south, the Alto de Juncal anticline can be seen. This large structure affects the Early Cretaceous volcanic upper plate of the Aconcagua fold-and-thrust belt. Large-scale folds are typical of this region (Fig. 33), but the intensity of folding decreases rapidly toward the west. Stop 5-8: Los Caracoles The Juncal Formation intruded by sigmoidal dikes and sills can be seen to the west at this stop. These dikes of dacitic composition were emplaced during Miocene thrusting (Godoy et al., 1999). Note the typical U shape of the glacial valley of the upper Río Aconcagua. Stop 5-9: Río Aconcagua Valley This stop shows a Miocene granodiorite emplaced into the Tertiary Abanico Formation (Charrier et al., 2005). The narrow gorge at Salto del Soldado is the result of a massive slide from

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the northeast. The stop provides an impression of the thick pile of andesitic flows, pyroclastic breccias, and tuffs that constitute the Abanico Formation that is mainly Oligocene in age. This unit has been deposited in a volcanotectonic rift episode (Godoy et al., 1999).

Stop 5-13: Coast between Concón and Valparaíso

Stop 5-10: San Felipe Quarry

REFERENCES CITED

Several quarries exploit algal (lagoonal?) limestones that are interfingered with the conglomerates of the Early Cretaceous Las Chilcas Formation. These limestones have typical palynomorphs of Neocomian age that can be correlated with the marine Neocomian limestones of the retroarc basin (Arévalo, 1992). A reverse fault and a related drag fold affect the sedimentary strata of Las Chilcas Formation, which is intruded by a 96 Ma granite. This unit is probably an age-equivalent of the Pelambres Complex. The Cordillera de la Costa The basement of the Cordillera de la Costa is formed by a displaced terrane (Pichidangui terrane). It was probably emplaced by margin-parallel, strike-slip movements in the Lower Jurassic. The basement is covered by a thick, autochthonous, east-dipping Mesozoic volcano-sedimentary terrane, and intruded by Jurassic and Cretaceous batholiths (Gana and Wall, 1997; Gana and Zentilli, 2000). The Central valley that separates the Cordillera de la Costa from the Cordillera Principal in much of Chile is absent at this latitude (33°S) and over the flat-slab segment to the north. Together, with the lack of active volcanism, these morphological attributes are typical of the modern “flat-slab” region of the Chile-Argentine Andes which extends between 27° and 33°S. Stop 5-11: Ocoa Quarry East-dipping outcrops of andesites (Ocoa member) of the Veta Negra Formation (Lower Cretaceous). The local term “ocoite” is used for a rock type with phenocrysts of plagioclase exceeding 2 cm in length. Radiometric dating (Rb/Sr) yielded an age of 105 Ma for this unit (Rivano et al., 1985; Morata and Aguirre, 2003). The Veta Negra Formation is the main volcanic (Aptian-Albian) unit of the Early Cretaceous central Chile volcano-tectonic rift. Note the association of alteration minerals (epidote-chlorite-actinolite-zeolite-albite). Originally attributed to “burial metamorphism,” this assemblage may represent the effects of low-temperature hydrothermal fluids circulating in the Early Cretaceous extensional environment. Stop 5-12: Papudo Gneissic tonalites (157–170 Ma) of the Jurassic Coastal Batholith are the oldest of the eastward younging series of “Andean” batholiths. Paleomagnetic studies have found no evidence of large displacement for this or younger rocks in the Aconcagua region.

Late Carboniferous tonalites (299 ± 31 Ma, Rb/Sr) in a narrow belt intruded to the east by Jurassic granitoids. K/Ar ages have been reset to the Jurassic (Hervé et al., 2000).

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Thomas, W.A., and Astini, R.A., 1999, Simple-shear conjugate rift margins of the Argentine Precordillera and the Ouachita embayment of Laurentia: Geological Society of America Bulletin, v. 111, p. 1069–1079, doi: 10.1130/0016-7606(1999)111<1069:SSCRMO>2.3.CO;2. Thomas, W.A., and Astini, R.A., 2003, Ordovician accretion of the Argentine Precordillera terrane to Gondwana: A review: Journal of South American Earth Sciences, v. 16, p. 67–79, doi: 10.1016/S0895-9811(03)00019-1. Thomas, W.A., Tucker, R.D., and Astini, R.A., 2000, Rifting of the Argentine Precordillera from southern Laurentia: palinspastic restoration of basement provinces: Geological Society of America Abstracts with Programs, v. 32, p. A-505. Triep, E.G., 1987, La falla activada durante el sismo principal de Mendoza de 1985 e implicancias tectónicas: X° Congreso Geológico Argentino, Actas, v. 1, p. 199–202. Uliana, M.A., Biddle, K.T., and Cerdan, J., 1989, Mesozoic extension and the formation of Argentine sedimentary basins, in Tankard, A.J., and Balkwill, H.R., eds., Extensional tectonics and stratigraphy of the North Atlantic margins: American Association of Petroleum Geologists Memoir, v. 46, p. 599–614. Valencio, D.A., Vilas, J.F., and Pacca, I.G., 1983, The significance of the paleomagnetism of a sequence of red beds of the Middle and Upper sections of Paganzo Group (Argentina) and the correlation of Upper Paleozoic Lower Mesozoic rocks: Geophysical Journal of the Royal Astronomical Society, v. 51, p. 59–74. Vergés, J., Ramos, E., Seward, D., Busquets, P., and Colombo, F., 2001, Miocene sedimentary and tectonic evolution of the Andean Precordillera at 31°S, Argentina: Journal of South American Earth Sciences, v. 14, p. 735–750, doi: 10.1016/S0895-9811(01)00070-0. Vergés, J., Ramos, V.A., Bettini, F., Meigs, A., Cristallini, E., Cortés, J., and Dunai, T., 2002, Geometría y edad del anticlinal fallado de Cerro Salinas, Mendoza: 15° Congreso Geológico Argentino, Actas, v. 3, p. 290–295. von Huene, R., Corvalán, J., Flueh, E.R., Hinz, K., Korstgard, J., Ranero, C.R., Weinrebe, W. and the Condor scientists, 1997, Tectonic control of the subducting Juan Fernández Ridge on the Andean margin near Valparaíso, Chile: Tectonics, v. 16, p. 474–488, doi: 10.1029/96TC03703. Vujovich, G., and Kay, S.M., 1998, A Laurentian?: Grenville-age oceanic arc/ back-arc terrane in the Sierra de Pie de Palo, Western Sierras Pampeanas, Argentina, in Pankhurst, R., and Rapela, C.W., eds., Protomargin of Gondwana: Geological Society [London] Special Publication 142, p. 159–180. Wegener, A., 1929, The origin of continents and oceans: New York, Dover Publication (Translation of the Fourth Edition), 231 p. Wehrli, L., and Burckhardt, C., 1898, Rapport préliminaire sur une expédition géologique dans la Cordillère argentino-chilienne entre le 33° et 36° latitud sud: Museo de la Plata, Revista, v. 8, p. 373–388. Wortel, M.J.R., 1984, Spatial and temporal variations in the Andean subduction zone: Journal of the Geological Society, v. 141, p. 783–791, doi: 10.1144/ gsjgs.141.5.0783. Yañez, G., Ranero, G.R., von Huene, R., and Diaz, J., 2001, Magnetic anomaly interpretation across a segment of the southern central Andes (32–34°S): Implications on the role of the Juan Fernández ridge in the tectonic evolution of the margin during the Upper Tertiary: Journal of Geophysical Research, v. 106, p. 6325–6345, doi: 10.1029/2000JB900337. Yañez, G., Cembrano, J., Pardo, M., Ranero, G.R., and Sellés, D., 2002, The Challenger-Juan Fernández-Maipo major tectonic transition of the NazcaAndean subduction system at 33–34°S: Geodynamic evidence and implications: Journal of South American Earth Sciences, v. 15, p. 23–38, doi: 10.1016/S0895-9811(02)00004-4. Yrigoyen, M.R., 1976, Observaciones geológicas alrededor del Aconcagua: I° Congreso Geológico Chileno, Actas, v. 1, p. 169–190. Yrigoyen, M.R., 1979, Cordillera Principal, in Turner, J.C.M., ed., Segundo simposio de geología regional Argentina: Córdoba, Academia Nacional de Ciencias, v. 1, p. 651–694. Zapata, T.R., and Allmendinger, R.W., 1996, The thrust front zone of the Precordillera thrust belt, Argentina: A thick-skinned triangle zone: Bulletin of the American Association of Petroleum Geologists, v. 80, p. 359–381. Zeil, W., 1979, The Andes: A geological review: Gebrúder Borntaeger, 260 p. Zeil, W., 1981, Volcanism and geodynamics at the turn of the Paleozoic to the Mesozoic in the Central and Southern Andes: Stuttgart, Zentralblatt für Geologie und Paläontologie, Teil 1–1981, p. 298–318. MANUSCRIPT ACCEPTED BY THE SOCIETY 10 JANUARY 2008 Printed in the USA

The Geological Society of America Field Guide 13 2008

Field trip guide: Neogene evolution of the central Andean Puna plateau and southern Central Volcanic Zone Suzanne Mahlburg Kay* Department of Earth and Atmospheric Sciences, Snee Hall, Cornell University, Ithaca, New York 14853, USA Beatriz Coira* Instituto de Geología y Minería, Universidad Nacional de Jujuy, Casilla de Correo 258 (4600), San Salvador de Jujuy, Argentina Constantino Mpodozis* Antofagasta Minerals, Ahumada 11, Santiago, Chile

ABSTRACT This seven-day field trip is designed to examine the distinctive magmatic, structural, and sedimentological features of the late Oligocene to Recent evolution of the southern part of the high central Andean Puna plateau and the southern Central Volcanic Zone magmatic arc. The stops for Days 1–5 between 23° and 27°S latitude in Argentina emphasize the distinctive magmatic and structural features of the Puna region, which comprises the southern half of the central Andean Puna–Altiplano plateau. Differences between the northern and southern Puna are highlighted. Among the features to be observed are (1) giant Miocene ignimbrites of the northern Puna, (2) distinctive normal and strike-slip faults and associated shoshonitic lavas of the central Puna, (3) the intraplate and calc-alkaline lavas of the southern Puna, (4) the silicic calderas of the southern Puna, and (5) internally drained salar basins. The stops for Days 6 and 7 between 26.5° and 27.5°S latitude in Chile present a view of the Miocene to Recent frontal arc region on the western side of the plateau. The stops particularly highlight the late Miocene to Pliocene displacement of the magmatic arc front from the Maricunga Belt on the western edge of the plateau to its present position in the Central Volcanic Zone. Evidence for the timing of plateau uplift, changes in the angle of the underlying subduction zone, delamination of the underlying continental crust and mantle lithosphere, and forearc subduction erosion are examined throughout the course of the trip. DISCLAIMER: This is not a geologic field guide to be used in the traditional sense of a detailed road log, but rather a survey of the southern Puna region that can be done with the use of Google Earth. Latitude and longitude coordinates are given for all stops in WGS84 (world geodetic system) coordinates that can be used with Google Earth (downloadable from the Web) or any other georeferenced imagery. Driving *E-mails: [email protected]; [email protected]; [email protected] Kay, S.M., Coira, B., and Mpodozis, C., 2008, Field trip guide: Neogene evolution of the central Andean Puna plateau and southern Central Volcanic Zone, in Kay, S.M., and Ramos, V.A., eds., Field trip guides to the Backbone of the Americas in the southern and central Andes: Ridge collision, shallow subduction, and plateau uplift: Geological Society of America Field Guide 13, p. 117–181, doi: 10.1130/2008.0013(05). For permission to copy, contact [email protected]. ©2008 The Geological Society of America. All rights reserved.

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Kay et al. instructions are given where access is possible by paved or high-quality unpaved roads that are present on road maps that are generally available in the region. Due to the lack of information on available road maps, changing driving conditions, and the need for caution in accessing the sites, precise instructions are not given for others. A significant number of these stops are on primitive, unmaintained roads or tracks that require a serious four-wheel drive vehicle, an experienced off-road driver, and supporting equipment (winch, spare tires, jack, etc.) to navigate safely. Many of the stops are in the Atacama Desert, where there is almost no water. The majority of the stops are at high elevations—most are between 3500 and 4500 m (~11,500–14,750 ft); all are over 2000 m (6500 ft). Attention needs to be paid to the potential for altitude sickness (called the Puna in Argentina). There are no cell phone towers, no service stations, or towing facilities in most of the region to bail you out! The availability of fuel can be questionable at times. Contact a knowledgeable guide to the region before attempting to use this field guide in the more isolated parts of the area. Keywords: Plateau uplift, delamination, ignimbrites, slab dip, arc migration, Central Andes.

INTRODUCTION TO THE FIELD AREA Overview of the Central Andean Puna-Altiplano Plateau The central Andes are known for a number of special features whose origins have been long debated. The most prominent geographic feature is the extensive Puna-Altiplano plateau (Fig. O.1), which after the Tibetan plateau is the world’s highest (average elevation 3700 m) and largest (700-km–long and 200-km–wide) plateau. Unlike the Tibetan plateau, which has little volcanic cover, the Puna-Altiplano is covered by an extensive array of Neogene volcanic centers that extend in chains to the eastern side of the plateau. The plateau is underlain by thick continental crust (>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

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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.

S

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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 <7 Ma mafic volcanic rocks (e.g., Kay et al., 1994a, 1999; Risse et al., 2008) and glassy andesitic to dacitic lava flows (e.g., Kay et al., 1999, 2005). Contrasts in the Magmatic and Structural History of the Northern and Southern Puna The northern and southern Puna are characterized by distinctive late Oligocene to Recent magmatic, structural, and sedimentological histories that are highlighted below and are the focus of the stops on Days 1–5. The magmatic front west of the southernmost Puna in Chile is the focus of Days 6 and 7. The main late Oligocene to Recent magmatic features of the northern and southern Puna are shown in Figures O.6 to O.8 and related to general tectonic models for transects in the northern (~21° to 24°S) and southern (~26° to 28°S) Puna in Figures O.9 and O.10. More complete discussions can be found in Coira et al. (1993), Allmendinger et al. (1997), Kay et al. (1999), Oncken et al. (2006), and references therein. Additional information on the magmatic arc front in the southernmost Puna can be found in Mpodozis et al. (1995, 1996) and Kay et al. (1994a, 2006).

Figure O.4. Crustal thicknesses from Yuan et al. (2002) based on critical reflections and P to S receiver functions averaged over one-degree grids on left; crustal thicknesses from McGlashan et al. (2008) based on seismic data at red points on right.

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Figure O.5. Interpreted lithospheric-scale profile based on depth-migrated seismic reflection data from ANCORP (2003).

mafic flows

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Figure O.6. Generalized distribution of Oligocene to Recent magmatic centers on the Puna plateau. Circles are silicic and ignimbrite eruptions; squares are mafic lavas and stratovolcanic centers; numbers are ages in Ma. CVZ—Central Volcanic Zone. Figure is modified from Kay et al. (1999).

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Figure O.7. Generalized distribution of Oligocene to Recent magmatic centers in the southern Puna (e.g., Kay et al., 2005). Red indicates silicic centers. Dotted and dashed lines show eastward migration of the arc front from 20 to 3 Ma. CVZ—Central Volcanic Zone.

Northern Puna History (Fig. O.6) The late Oligocene to middle Miocene (28–17–16 Ma) history of the northern Puna is characterized by a general gap in magmatic activity, contractional deformation, and foreland sedimentary basin development from the arc to the backarc of the modern plateau. This situation changes in the middle Miocene at 15–12 Ma as small generally andesitic porphyritic stocks and domes erupted in the far backarc (e.g., Caffe et al., 2002). In the late Miocene, the picture changed dramatically as large ignimbrite centers erupted across the plateau (de Silva, 1989; Coira et al., 1993; Kay et al., 1999) and contractional deformation ceased on the plateau (e.g., Gubbels et al., 1993). The eruptions include those of the ~150 km3 Granada center at

ca. 10 Ma (Caffe et al., 2007), the ~1400 km3 Vilama center at ca. 8.5 Ma (Soler et al., 2007), the ~650 km3 Panizos center at ca. 6.7 Ma (Ort, 1993; Ort et al., 1996), and the ~650 km3 Coranzulí center at ca. 6.7–6.4 Ma (Seggiaro and Aniel, 1989; Seggiaro, 1994) (volumes in dense rock equivalents [DRE]). After ca. 10 Ma, the main locus of foreland basin formation and thrusting shifted to the Subandean Belt to the east where thrusting began at ca. 9–8.5 Ma in the west and at ca. 6.9 Ma to the east (Echavarria et al., 2003). Major surface uplift at this time based on paleobotany (Gregory-Wodzicki, 2000) is supported by oxygen isotopic analyses of carbonates (Garzione et al., 2006). Large ignimbrite eruptions continued in the latest Miocene to Pliocene with the centers concentrated far-

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Figure O.8. Map showing migration of arc front from the Maricunga Belt to the Central Volcanic Zone (CVZ) arc, generalized distribution of late Pliocene to Pleistocene calc-alkaline, intraplate-like and shoshonitic mafic flows, and the Cerro Galan ignimbrite. Map is modified from Kay et al. (1999).

ther to the west. The eruptions include those of the >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 <3.8 Ma: In the most recent stage, andesitic and dacitic magmatism has been essentially confined to the arc front region by the steeper subduction zone, and shortening has continued in the Subandean Belt. The eruption of small backarc shoshonitic lavas (Kay et al., 1994b) and seismic images (e.g., Beck and Zandt, 2002; Schurr et al., 2006) are in agreement with removal of pieces of lithosphere beneath the Eastern Cordillera.

Neogene evolution: central Andean Puna Plateau and southern-central volcanic zone Maricunga front 22 to 7 Ma

backarc centers 12 to 7 Ma

0 km 100 200

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300

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Antofalla region mafic flows Cerro Galán

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~ 7 to 3 Ma

300 0

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Figure O.10. Lithospheric-scale sections to explain the magmatic and deformational history of the southern Puna. Lower section is modified from Kay et al. (1994a). CVZ—Central Volcanic Zone.

Southern Puna: Shallowing to Steepening Subduction and Crust and Mantle Delamination (Fig. O.10) The model for the southern Puna in Figure O.10 involves moderate middle Miocene shallowing of the subduction zone, late Miocene steepening, and late Miocene to Pliocene lower crustal and lithospheric delamination coincident with frontal arc migration. The delamination model is modified from Kay et al. (1994a, 1999) and the entire model is presented in Kay et al., 2005. Stage 1 at ca. 26–20 Ma: The concentration of latest Oligocene and early Miocene andesitic to dacitic volcanism in the arc is consistent with a relatively steeply dipping slab (Fig. O.8). A noncontractional setting with local extension fits with eruption of voluminous dacitic-rhyolitic domes and ignimbrites in the arc and the Segerstrom mafic lavas in the backarc. Stage 2 at 19–16 Ma: A change to a compressional stress regime at this time is consistent with contractional deformation in the backarc in Chile (e.g., Mpodozis and Clavero, 2002) and in the Puna to the east (e.g., Kraemer et al., 1999). The initiation of Miocene contraction along much of the Andean chain (e.g., Kay and Mpodozis, 2002; Oncken et al., 2006) at this time has been suggested to be due to the accelerated westward drift of South America over the Nazca plate (Kay and Copeland, 2006; Oncken et al., 2006) in the manner suggested by Silver et al. (1998). Stage 3 at 15–8 Ma: Shallowing of the subduction zone during this period can explain the eastward broadening of the

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magmatic arc, the arc-like geochemistry (high field strength element depletion; Kay et al., 2005) of the backarc volcanic rocks and backarc contractional deformation. A period of shallowing fits with subduction of the Juan Fernandez ridge on the Nazca plate playing a role in the shallowing because the reconstruction of Yañéz et al. (2001) shows the ridge approaching and subducting under the region from ca. 14–8 Ma. Shallowing needs to have been less pronounced than under the northern Puna because no temporal magmatic gap is observed. Stage 4 at 7–4 Ma: A major change in magmatic and deformational style at ca. 7 Ma is consistent with steepening of the subducting slab and delamination of the lower crust and underlying mantle lithosphere. Steepening of the slab is again consistent with the Yañéz et al. (2001) model that shows the Juan Fernandez Ridge subducting farther to the south by this time. A regional change in the deformational style to a complex mix of normal, strike-slip, and contractional faults (Marrett et al., 1994), the eruption of mafic lavas along faults, and uplift of the plateau (Alonso et al., 2006) fit with delamination of gravitationally unstable lower crust into the mantle wedge at this time (Kay and Kay, 1993; Kay et al., 1994a, 1999). The resulting influx of asthenospheric mantle into the thickening mantle wedge can explain mantle melting leading to basaltic volcanism and the widespread crustal melting producing the early Cerro Galan ignimbrite eruptions. This is also the time of the extinction of the Maricunga Belt arc and widespread volcanic activity between the Maricunga Belt, and the future Central Volcanic Zone arc front (Figs. O.7 and O.8; Mpodozis et al., 1995, 1996; Kay et al., 1994b; Kay and Mpodozis, 2002). Migration of the arc front is consistent with a major pulse of forearc subduction erosion (Kay and Mpodozis, 2002; Kay, 2006). Stage 5 after 3 Ma: By 3 Ma, a large amount of crustal melting following delamination resulted in the major Cerro Galán ignimbrite eruption at ca. 2.2 Ma. A thick mantle wedge fits with eruption of intraplate-like lavas and the geophysical evidence for a thin mantle lithosphere (Sn attenuation; see Whitman et al., 1992, 1996). FIELD TRIP LOG PART 1: DAYS 1 TO 5— THE ARGENTINE PUNA Beatriz Coira and Suzanne Mahlburg Kay The objectives of Days 1–5 are to examine the latest Oligocene to Recent history of the Puna plateau and to briefly view the Paleozoic basement. The route of the trip, which generally progresses from north to south, is shown in Figure 1.1. Day 1 features the Eastern Cordillera and the northern Puna; Day 2 features the transitional zone between the northern and southern Puna; Days 3, 4, and the first part of Day 5 feature the southern Puna, and the last part of Day 5 features the transition to the northern Sierras Pampeanas. A general overview of these regions is presented in the Introduction to the field area.

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S

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Day 2

Day 3

Days 6&7

Day 4

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Figure 1.1. Map of northwestern Argentina showing the field trip route in Argentina. Black lines with arrows indicate the route, and white dots show end points and overnight lodging sites for each day. Days 2–4 all terminate in the village of Antofagasta de la Sierra.

DAY 1—NORTHERN PUNA

Stop 1-1: Arroyo del Medio Alluvial Cone

The objective of Day 1 is to examine a transect from the Eastern Cordillera onto the northern Puna. The route starts in the city of Jujuy, goes to the north and then to the west ascending through the Eastern Cordillera fold-and-thrust belt onto the Puna (see Figs. 1.2 to 1.4).

Directions: Take Argentine Route 9 north from the city of Jujuy in the direction of Humahuaca for ~36 km. Stop along road near 23°56′50″S; 65°27′30″W; 2100 m asl. The large alluvial cone seen here is principally composed of the series of debris flows that dammed the valley of the Rió

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W 66°00′ W

Geologic Map Province of Jujuy

Panizos Ignimbrite Granada Ignimbrite Tmi

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Figure 1.2. Geologic map of the Argentine province of Jujuy showing an overview of the stops for Day 1 (numbers in white) and the morning of Day 2 (numbers in black) and the locations of the city of San Salvador de Jujuy and the villages of Purmamarca, Susques, and Huancar. Principal geologic units are the Precambrian-Cambrian Puncoviscana Formation (PCpv; mainly sedimentary rocks), the Cambrian Meson Formation (Cs; principally quartzite), Ordovician sedimentary rocks (Os), Ordovician volcanic and plutonic rocks (mainly silicic; Om), Jurassic and Cretaceous granites (JKg), Cretaceous rift deposits (K), Miocene sedimentary rocks (Tms), principal late Miocene and Pliocene ignimbrites and volcanic rocks (Tmi and Tpi, respectively), and Quaternary clastic sediments (Q) and evaporites (Qe).

Grande in 1945. The narrowing of the valley at this point is a result of these debris flows. The large size of the flows reflects both the active seismicity of the region and the coincidence of this sector of the river valley with the line of maximum storm intensity during the rainy season (January and February). The voluminous Holocene to Recent clastic sediments along the Quebrada de Humahuaca to the north further reflects the crustal seismicity that demonstrates the ongoing tectonic activity in the region. A seismic array deployed by the PANDA group (Cahill et al., 1992) along the eastern edge of the Puna, the Eastern Cordillera, and into the Santa Barbara structural belt elucidated the locations of the crustal earthquakes associated with contractional structures along the eastern deformational front. A peak of seismicity at a depth of 20–25 km and fault plane solutions

for these earthquakes are in accord with a mid-crustal detachment on which significant east-west late Cenozoic foreland shortening has occurred (Cahill et al., 1992). Stop 1-2: View of Typical Eastern Cordillera Structures Just East of Purmamarca Directions: Continue north on paved Route 9 and turn left (west) on paved Route 52 toward the village of Purmamarca. Stop is at a pull-out on the right side of the road ~3 km after the turn and before entering Purmamarca. Stop is at 23°44′50″S; 65°29′39″W; 2313 m asl. The Eastern Cordillera is a thrust belt with a typical piggy-back structural style that shows both west and east vergence. Balanced

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65°30′W Ks

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MPs

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Figure 1.3. Portion of the Ciudad de Liberator General San Martín Servicio Geológico Minero Argentino (SEGEMAR) geologic map (González et al., 2000) originally mapped at a scale of 1:250,000 showing the geology at Stops 1-1 to 1-5. Map illustrates the imbricated faulted structure of the Eastern Cordillera. Principal units are the Precambrian-Cambrian Puncoviscana Formation (PCpv; mainly sedimentary rocks), Cambrian granitoids (Cg), the Cambrian Meson Formation (Cs; principally quartzite), the Ordovician Santa Victoria group (sedimentary rocks; Os), Jurassic and Cretaceous granites (JKg), Cretaceous rift deposits (Santa Pirgua, Balbuena, and Santa Barbara subgroups; green colors—Ks), Miocene and Pliocene sedimentary (MPs), and various Quaternary to Recent sediments (Qs).

cross sections indicate ~60 km of shortening across the Eastern Cordillera. Cenozoic deformation and uplift of the Eastern Cordillera began at ca. 40 Ma and were largely over by 10 Ma (Müller et al., 2002; Ege et al., 2007). Some deformation is still occurring as shown by earthquake activity and evidence for very recent thrust and strike-slip deformation. An array of west- and eastverging thrusts involves late Precambrian, Lower Paleozoic, and Cretaceous to Paleogene stratigraphic units (Figs. 1.3 and 1.5). At this stop, imbricated slices of Cambrian and Cretaceous-Paleocene rocks can be seen thrust over the sedimentary sequences of the Precambrian–Lower Cambrian Puncoviscana Formation. The complex structures in this region are interpreted as having formed

from a combination of west-verging compressional thrusts in the Late Ordovician (Ocloyic Event), extensional rift faults in the Cretaceous to Paleogene, and east-verging contractional structures in the Miocene to Pliocene. The multi-episodic deformation resulted in a wide variation in vergence and partial inversion of extensional structures leaving younger over older fault relationships. Purmamarca Village This is a tourist spot for lunch and a great place to buy local merchandise typical of the Puna plateau region. Among the attractions is the Santa Rosa de Purmamarca chapel, which

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

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6 23°30′S

5 4

Salinas Grandes

3 2

Purmamarca 66°W

1 24°S

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66°30′W

66°W

Figure 1.4. Thematic Mapper satellite image showing the route from Jujuy to the west side of the Salinas Grandes and the locations of Stops 1-1 to 1-6.

east

west 4000 m 0m –5000 m –10000 m

Ticara Profile –15000 m

Figure 1.5. West to east structural cross section near 23°45′S across the Eastern Cordillera from the Servicio Geológico Minero Argentino (SEGEMAR) geologic map of Ciudad de Liberator General San Margin (González et al., 2000). The section shows the general structure of the Eastern Cordillera at the latitude of Stops 1-2 to 1-5.

is a Jesuit church that was constructed between 1648 and 1779. Paintings of the seventeenth century Cuzco school exhibited in the chapel depict the life of Santa Rosa de Lima. Other paintings are entitled La Piedad and La Inmaculada. At the chapel entrance stands a five-century–old algarrobo tree. The town preserves a Spanish colonial style with a strong indigenous and Argentine overprint. A native market place is located in the main square.

Stop 1-3: View of Alluvial Deposits Looking East from the Cuesta de Lipan Directions: Continue west on Route 52 up the eastern side of the plateau. Stop at the Cuesta de Lipan near 23°40′19″S; 65°36′13″W; 3390 m asl. View is to the east. The outstanding sequence of highly colored PliocenePleistocene deposits of the Purmamarca Formation that formed

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under periglacial conditions can be viewed from this stop. The deposits are strongly dissected by the Purmamarca River marking the ongoing uplift of the Eastern Cordillera. A monadnock, which is an isolated remnant of a former erosional cycle, can be seen in a tributary valley. Stop 1-4: View of Cretaceous Rocks Affected by Cenozoic Deformation at El Angosto Directions: Continue west on Route 52 climbing the eastern side of the plateau, passing the summit, and start descending the switchbacks to the west. Stop is at El Angosto near 23°42′W; 65°41′W; 3891 m asl. Structures associated with Cretaceous rifting events have had a profound influence on the style of Cenozoic thrusting and folding in the Eastern Cordillera. Many north- to northnortheast–trending normal faults that were active during the Cretaceous are oriented approximately perpendicular to the Cenozoic contractional direction. These former normal faults have been reactivated as reverse faults. At this stop, Andean age Miocene-Pliocene thrusts can be seen to cut CretaceousPaleocene extensional faults that affected the synrift Cretaceous Pirgua Subgroup. The post-rift deposits of the Paleogene Balbuena and Santa Barbara Subgroups are also involved in the thrusts. Thrusts that put sheets of Precambrian-Cambrian basement over younger units are visible at this stop. Stop 1-5: Salar de Salinas Grandes Directions: Continue west on Route 52 to Salinas Grandes. Stop at the tourist stop in the middle of the salar near 23°35′47″S; 65°52′58″W; 3429 m asl. Salars are a distinctive feature of the Puna. They are basically intermontane basins that have principally developed over the past 7–8 Ma as large volumes of continental evaporites accumulated during the uplift of the Puna. Evaporate deposition is controlled by geothermal activity related to volcanism, internal drainage, and climate. The Salinas Grandes salar at this stop hosts one of the major halite deposits of the Puna. Production is 10,000–20,000 tons per year depending on the weather. Borate deposits, which are mainly from ulexite [NaCaB5O6 (OH)6-5H2O], are exploited in the Tres Morros and Niño Muerto sectors of the salar. The Salinas Grandes is a symmetrical basin bounded by inverted Cretaceous high-angle normal faults. An eastsoutheast–trending seismic reflection profile between the Salinas Grandes and Laguna de Guayatayoc from Coutand et al. (2001) in Figure 1.6 shows three main units. Unit 1—Paleozoic basement units and the Early Cretaceous Tusaquillas pluton; Unit 2—Cambrian to Early Ordovician quartzites subjected to folding and faulting prior to Andean deformation (Gangui, 1998); and Unit 3—Mesozoic and Cenozoic strata separated from Units 1 and 2 by a major angular unconformity (Gangui, 1998). The lower units show evidence for Cretaceous

normal faults that were inverted during Andean compression (Acevedo and Bianucci, 1987). The Salinas Grandes depocenter lies between two major thrusts that were active during Cenozoic sedimentation. Stop 1-6: Ordovician Basement (Cobres Plutonic Complex: Churcal and Las Burras Granitoids) and Distal Miocene Ignimbrite Flows Directions: Continue west on Highway 52, crossing the Salinas Grandes, and stop on the north side of the road after entering the low ranges to the west. Stop is at ~23°25′21″S; 66°11′18″W; 3525 m asl. The deformed magmatic and sedimentary rocks seen here and along the field trip route to the west (Fig. 1.2) are representative of Ordovician sequences, which form an important component of the Puna basement. The exposures in the northern Puna are dominated by low- to medium-grade metamorphic rocks, syntectonic rhyolites, and rare granitoids. The exposures in the southern Puna are dominantly composed of higher grade metamorphic rocks and abundant granitoids. The dominance of Ordovician basement outcrops in the Puna contrasts with the situation in the Eastern Cordillera, where the principal exposures are Precambrian to Cambrian sedimentary sequences. The outcrops at Stop 6 (Figs. 1.2 and 1.7) are part of the syntectonic lower Ordovician Cobres Plutonic Complex (Kirschbaum et al., 2006), which is composed of granodiorite and monzogranite plutons. The plutons record a tectonomagmatic event at 476 Ma (Lork and Bahlburg, 1993; Haschke et al., 2005), which is the time of collision of the Laurentian-derived Precordillera (Cuyania) terrane with the Gondwana continent to the south (Coira et al., 1999). An initial low-temperature deformational event produced the intense folding and cleavage in the exposures in this region. Subsequent medium-temperature deformational features can be associated with pluton emplacement and metamorphism. The last event in this tectonic cycle was the intrusion of the post-tectonic Las Burras Granite (428 ± 17 Ma, Zappettini, 1989). The Ordovician basement and valleys along the field trip route are covered by the distal flows of the Las Termas ignimbrite (6.45 Ma; Seggiaro, 1994) that erupted from the Coranzulí center ignimbrite center (Fig. 1.8). Stop 1-7: Las Termas Ignimbrite North of the Village of Barrancas Directions: Turn right (north) on a prominent secondary dirt road near Stop 1-6 to the village of Barrancas. Stop 1-7 is 15 km northeast of Stop 1-6 at the well-exposed ignimbrite exposures near 23°19′50″S; 66° 05′28″W; 3624 m asl. After the stop, return to Route 52. The Las Termas ignimbrite (6.45 ± 0.15 Ma; Seggiaro, 1994) at this stop is the longest and youngest of the four ignimbrite sheets that erupted from the late Miocene Coranzulí caldera

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

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Figure 1.6. Interpreted time-migrated seismic line across the Salinas Grandes basin south of Stop 1-5. Vertical scales are in two-way travel time. See description of Stop 1-5. Figure modified from Coutand et al. (2001).

66°45′ W

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Laguna de Guayatayoc

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Barrancas Susques

1 Salar Olaroz

8 Pastos Chicos basin

6 Salinas Grandes

Figure 1.7. Thematic Mapper satellite image showing the stops between the Salinas Grandes and the village of Susques. The Coranzulí caldera-resurgent center seen from Stop 8 is well seen. This caldera is the source of the Las Termas ignimbrite at Stop 7 near the village of Barrancas and elsewhere on the route to Susques. The Pastos Chicos basin is marked by the deposits of Miocene Pastos Chicos Formation.

(Figs. 1.7 and 1.8). The Coranzulí caldera, which is some 35 km to the north (Fig. 1.7), is better seen at the next stop. The Las Termas ignimbrite is well exposed all along the Barrancas valley, where its distribution was controlled by preexisting relief. The Las Termas ignimbrite is a moderately to highly welded, crystal-

rich (plagioclase, sanidine, and biotite) dacite with a moderate amount of pumice. The flow unconformably overlies deformed sedimentary deposits of the middle Miocene Pastos Chicos Formation and puts a minimum age on the last stage of contractional deformation in this region of the Puna.

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68°00′ W

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LEGEND Calderas structural limit Lineament or regional fault COT (Calama-Olacapato El Toro fault zone)

Panizos Caldera

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CERRO GUACHA CALDERA

Vilama Caldera

Post collapse lava- dome 7.5-8.1 Ma Intracaldera Ignimbrite Lower unit, extracaldera Ignimbrite Upper unit, extracaldera 8.4-8.5 Ma

Cerro Guacha Caldera

Post-collapse lava-dome 2.1 Ma Upper Tara 3.4-3.8 Ma Puripicar Ignimbrite Totoral Lower- Tara Ignimbrite 5.6 Ma Pampa Guayaques Ignimbrite 5.9 Ma Kalina Tobas

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CORANZULI CALDERA Coranzulí Caldera LA PACANA CALDERA

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Post collapse lava - dome 6.1 Ma Upper cooling unit Cerro Panizos Lower cooling unit Ignimbrite 6.7 Ma

AGUAS CALIENTES CALDERA

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Aguas Calientes Caldera

Post collapse lava-dome 2-3.8 Ma Atana Ignimbrite 3.8-4.2 Ma Toconao Ignimbrite 4-5.3 Ma Abra de Gallo Ignimbrite (10-10.5 Ma) Tajamar and Chorrillos Ign.+Olacapato and Sta.R.P.Grandes Ign (10-15 Ma) Volcanic Complex A.Calientes (17.2 Ma)

Figure 1.8. Map shows the major calderas and ignimbritic units of the northern Puna (Coira et al., 1993; Coira, 1998). The four major ignimbrite flows associated with the Coranzulí center are as mapped by Seggiaro (1994).

Stop 1-8: View of Coranzulí Caldera–Resurgent Center and Ignimbrite Flows Directions: Continue west on Route 52, stop near 23°26′26″S; 66°17′29″W; 3852 m asl. The Coranzulí Caldera complex, seen to the north at this stop, is located at the intersection of major northeast- and northtrending regional faults (Fig. 1.8) that controlled the relative movement of the Rinconada and Cochinoca structural blocks. These faults, together with the northwest-striking Ramallo fault, define a transtensional system, which controlled the development of the collapse caldera (Seggiaro, 1994). On a regional scale, the northeast-trending faults define part of the Lipez-Coranzulí lineament that traverses the Puna. Volcanic centers aligned along this lineament form one of the distinctive transverse Puna volcanic chains. According to Riller et al. (2001), onset of activity on these faults at ca. 10 Ma marks an important temporal transition from a deformational regime controlled by dominantly vertical thickening to one controlled by orogen-parallel stretching as critical crustal thicknesses were achieved in the Puna. The Coranzulí Caldera complex erupted four discrete ignimbrite sheets between 6.8 and 6.5 Ma (Seggiaro et al., 1987). These ignimbrites unconformably overlie deformed Miocene sedimen-

tary sequences that fill the Pastos Chicos sedimentary-tectonic depression. The combined volume of the four ignimbrites is estimated to be over 650 km3 (Seggiaro et al., 1987). All four have rhyodacitic compositions, moderate amounts of pumice, high crystal contents, and moderate welding. The proximal and distal facies are similar. Ash-fall deposits are scarce. Ground surge deposits can be seen in some regions. The eruption of the 6.5 Ma Las Termas ignimbrite can be related to the collapse caldera of the Coranzuli center, which has a diameter of 5 km. Eruption was followed by a magmatic resurgence that is represented by the late Miocene Coranzulí lava dome. The dome fills and partially obscures the caldera depression. DAY 2—TRANSITION ZONE BETWEEN THE NORTHERN AND SOUTHERN PUNA The objective of Day 2 is to examine the Neogene evolution of the transition zone between the northern and southern Puna. The route will cross the El Toro lineament, which is the boundary between the southern and northern Puna (Alonso et al., 1984b). South of 24°S, late Neogene strike-slip and extensional faults, small discontinuous and diachronous sedimentary basins, and mafic lavas appear.

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

Directions: 23°25′16″S; 66°22′58″W; 3660 m asl. After the stop, take a secondary dirt road just east of the Pastos Chicos Hostería to the south in the direction of the villages of Huancar and Sey (see map in introduction to Days 1–5). This stop provides a view of clastic sedimentary sequences in the Pastos Chicos Formation that lie unconformably above deformed lower Ordovician sedimentary basement rocks. A tuff in the Pastos Chicos Formation has been dated at 9.5 ± 0.3 Ma (Schwab, 1973). At this stop, the Pastos Chicos Formation is unconformably covered by the undeformed 6.4 Ma Las Termas ignimbrite from the Coranzulí center. South of Susques and the Rio Sijes, the Pastos Chicos Formation unconformably overlies deformed sediments in the Miocene Trinchera Formation (Schwab, 1973) that were deposited unconformably on Ordovician and Cretaceous units. The Trinchera Formation consists of clastic and calcareous beds intercalated with tuffs and ignimbrites that have yielded a K/Ar age of 10.8 ± 0.3 Ma (Schwab and Lippolt, 1976). On the flank of the Salar de Olaroz-Chachari to the west (see Figs. 1.7 and 2.1), Ordovician sedimentary rocks are thrust over conglomerates in the Pastos Chicos Formation. To the north, undeformed ca. 10 Ma Granada and ca. 8–7 Ma Vilama ignimbrites (Fig. 1.8) overlie deformed Miocene clastic and pyroclastic rocks (Soler et al., 2007). The regional relations show that late Miocene shortening was going on in the Susques region before the eruption of the Las Termas ignimbrite at 6.4 Ma and that deformation had ceased by 9–10 Ma. Similar field relations have been used to argue that contractional deformation had ceased in the northern Puna and southern Altiplano and shifted into the eastern foreland by ca. 10–9 Ma (Gubbels et al., 1993). Optional Stop near Village of Huancar: Lower Ordovician Sedimentary and Volcanic Rocks under Deformed Late Miocene Sediments of Pastos Chicos Basin The deformed Lower Ordovician turbidities, synsedimentary dacitic lavas and hyaloclastites, and sparse mafic lavas that form the basement under the Miocene sequences (Coira, 1996; Coira et al., 1999) are well exposed near the village of Huancar. As elsewhere in the northern Puna, the Ordovician sequence is characterized by low-grade metamorphism. The overlying late Miocene clastic, evaporitic, and pyroclastic sequences of the Miocene Trinchera and Pastos Chicos Formations record the evolution of the Pastos Chicos intermontane basin between 10.8 and 8.9 Ma (Schwab and Lippolt, 1976). Stop 2-2: Eruption Sequence and Origin of Tuzgle Volcano Directions: Continue on secondary dirt road. Stop is 91 km south of the village of Susques at 23°50′56″S; 66°27′45″W; 3900 m asl. See Figures 2.2 and 2.3. From this stop, the Tuzgle volcanic complex can be seen to the south. Tuzgle volcano is on the northern side of the El Toro

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Figure 2.1. Generalized geologic map of the east-central Puna from Coira and Kay (1993) showing Tuzgle volcano, the San Gerónimo and Negro de Chorrillos shoshonitic centers, and the Quevar volcano. The northwest-trending fault north of Chorrillos is the El Toro fault.

lineament, which is a first-order, northwest-trending transtensional fault system that crosses the plateau (see Fig. 2.1). The small San Gerónimo and Negro de Chorrillos shoshonitic volcanoes seen at Stop 2-7 are along the southern side of the lineament (Figs. 2.4 and 2.5). On a regional scale, the Tuzgle volcano is located near the eastern edge of the Puna plateau, some 275 km east of the main front of the Central Volcanic Zone and some 200 km above the Wadati-Benioff zone. The underlying subducting Nazca plate marks the northern end of the transition zone where the slab begins to shallow into the Chilean flat slab south of 28°S (Cahill and Isacks, 1992; Figs. O.1 and O.2). The Tuzgle eruptive sequence in Figure 2.2 is discussed in Coira and Kay (1993). Volcanic activity (see Fig. 2.2) began at 0.5 ± 0.2 Ma (Aquater, 1980) with the eruption of ~0.5 km3 of rhyodacitic ignimbrite that flowed to the north. The ignimbrite was followed by the 0.3 ± 0.1 Ma Old Complex dacitic lava dome complex that has a total volume of ~3.5 km3. Next, the Old Complex

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Figure 2.2. Top: Geologic map of Tuzgle volcano and surrounding region from Coira and Kay (1993) and Coira (1995). Bottom: Thematic Mapper satellite image of Tuzgle volcano showing the location of Stops 2-2 through 2-5 and expression of the volcanic units.

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

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was partially covered by the pre-platform andesitic lava flows that erupted from a new crater. This new crater was in turn partially filled by the Platform mafic andesite flows. These early-stage Tuzgle volcanic rocks were then cut by northwest- and east-trending faults that controlled the emplacement of the Post-platform and Young Tuzgle mafic andesite lava flows (Coira and Paris, 1981). The total volume of the Pre-platform and Young lavas is ~0.5 km3. The Tuzgle region is an active geothermal area (Coira, 1995) with the Tuzgle and Agua Caliente hot springs having temperatures of 40 °C to 56 °C and the Planta Mina hot springs having a temperature of 21 °C. Geothermal reservoir temperatures are 132 °C to 142 °C. Geochemical studies indicate mixing of deep and shallow water. Petrologic and geochemical data show that the Tuzgle volcanic rocks are notably diverse and require variable thermal conditions and diverse mantle and crustal components to explain

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Figure 2.3. Top: Interpreted P-wave tomographic section at 24.2°S latitude from Schurr et al. (2003). The areas in red beneath the Central Volcanic Zone (CVZ) in the western Cordillera (WC) and Tuzgle volcano are interpreted as zones where H2O has been released from the subducting oceanic lithosphere producing melts. The white dots in the slab are earthquakes attributed to dehydration reactions in the slab. The arrows indicate melt ascent paths. Bottom: Lithospheric section from Coira and Kay (1993) showing pathways (dashed lines) for intraplate-like and transitional shoshonitic magmas at Tuzgle volcano and shoshonitic lavas at San Geronimo and Negro de Chorrillos. Shape of the subducting plate is from Cahill and Isacks (1992), and crustal décollement at ~20–25 km is from Cahill et al. (1992). The base of the continental lithosphere is shown as being delaminated with blocks from the west containing crust. Mantlederived magmas in the asthenospheric wedge above the subducting plate are shown as being contaminated by melts of slab component-enriched lithospheric blocks before ascending into the overlying lithosphere and crust where they are further contaminated and fractionate. Magmas erupted to produce the rhyodacitic ignimbrites accumulated and crystallized at the décollement at ~20 km. The subsequent magmas evolved by mixing of rhyodacitic and more mafic magmas at the same level. The post-platform and Young Flow lavas are the lowest percentage mantle melts with the greatest lithospheric contamination among the Tuzgle magmas. The Negro de Chorrillos and San Geronimo shoshonites were produced by lower percentage mantle melts with more lithospheric contamination.

their origin (Coira and Kay, 1993). One general group includes the Tuzgle ignimbrite, Old Complex, and pre-platform units, which all have intraplate-like high field strength element (HFSE) signatures and rare earth element (REE) patterns with La/Yb <30. A second group includes the post-platform and Young andesitic flows, which all have arc-like HFSE anomalies and very steep REE patterns with La/Yb >35. A third group in the region includes the San Gerónimo and Chorrillos shoshonites to be seen at Stop 2-7. All of these magmas have isotopic and trace element evidence for upper crustal contamination. Major element and 87 Sr/86Sr ratio similarities between plagioclase (~AN33) xenocrysts in plagioclase-phenocryst–free mafic lavas and ignimbrite phenocrysts suggest that the plagioclase and quartz xenocrysts in the mafic lavas were introduced by mixing of variable composition mafic magmas with the rhyodacitic like those which produced the ignimbrites magmas (Kay et al., 1994a).

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Figure 2.4. Portions of the Servicio Geológico Minero Argentino (SEGEMAR) San Antonio de Los Cobres (Blasco et al., 1996) and Cachi (Hongn and Seggiaro, 2001) geologic maps compiled on a scale of 1:250,000. Stops 2-3 to 2-15 are indicated. Labeled units are PCmet—Precambrian metamorphic complex; PCgr—Precambrian granitoids; PcCpv—Precambrian to Cambrian Puncoviscana Formation; Ogr—Ordovician granitoids; C—Cambrian quartzites; O—Ordovician; mainly sedimentary rocks; Ofe—Ordovician Faja Eruptiva; K—Cretaceous rift and post-rift deposits; Teg—Tertiary Eocene Geste Formation; Tmp—Tertiary Miocene Pozuelos Formation; Tms—Tertiary Miocene Sijes Formation; Tmc— Tertiary Miocene Catal Formation; Tmv—Tertiary Miocene volcanics; Mios&v—Miocene sedimentary and volcanic rocks; Tps—Tertiary Pliocene Siguel Formation; Tpv—Tertiary Pliocene volcanics; Pleis—Pleistocene sediments; Qv—Quaternary volcanics; and Q—Quaternary sediments.

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

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Figure 2.5. Thematic Mapper satellite image of the monogenic San Jerónimo and Negro de Chorrillos shoshonitic centers showing their relation to Chorrillos and Incachule faults in the northwest-trending Calama–Olacapato–El Toro fault system (megatraza del Toro). Faults based on those in Marrett and Strecker (2000).

Figure 2.3 from Coira and Kay (1993) shows a model to account for these diverse volcanic rocks. The model involves (1) melting in the asthenospheric mantle above the subducting slab; (2) variable contamination of the melts by delaminated slab-modified lithosphere in the asthenosphere and by in situ continental lithosphere and crust; and (3) magma accumulation, fractionation, and mixing of diverse mafic magmas with the rhyodacite that produced the ignimbrite at décollement depths near 20 km before eruption. The most shoshonitic magmas require the lowest percentage mantle melts and the most slab-modified lithospheric contamination. The characteristics of the Tuzgle magmas are consistent with the underlying mantle being in a transitional setting from a thick continental lithosphere in the incipient stages of delamination in the Altiplano to the north to a thinner post-delamination mantle wedge producing calc-alkaline and intraplate magmas in the southern Puna (Kay et al., 1994a, 1999). Hoke et al. (1994) and Myers et al. (1998) argued for incipient delamination under the Altiplano based on mantle 3He/4He anomalies and seismic data, respectively. In the Tuzgle region, Schurr et al. (2003) use a tomographic model (Fig. 2.3) to argue for melting or fluids ascending from an earthquake cluster at 200 km depth into the crust below Tuzgle. Schurr et al. (2006) interpret a more recent tomographic image below Tuzgle as showing blocks of continental lithosphere in the process of being delaminated, melting in the lower crust, and melt accumulation at ~20 km in accord with the model of Coira and Kay (1993) in Figure 2.3. Stop 2-3: Exposure of the Tuzgle Ignimbrite Directions: Continue south on secondary dirt road. Stop is 2 km south of the village of Sey at 23°57′28″S; 66°29′35″W; 4040 m asl. The Tuzgle ignimbrite is well seen on the west side of the road at this stop. This ignimbrite, which was the first magma to

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erupt from the Tuzgle center, flowed to the north over Miocene sedimentary and volcanic rocks and the Ordovician volcanic and sedimentary sequences of the Faja Eruptiva Oriental Formation. The flow can be seen to be moderately crystal rich and to have a low degree of welding. The basal part of the ignimbrite contains sparse lithic fragments that are mainly derived from the underlying Ordovician basement. Pumice fragments constitute 10%–20% of the middle to upper sections of the ignimbrite. Stop 2-4: Exposure of the Tuzgle Pre-Platform Andesite Flows Directions: Continue south on dirt road; stop at 24°05′13″S; 66°30′47″W; 4461 m asl. The blocky lava of the Tuzgle pre-platform andesite unit is well seen at this stop. The phenocrysts in the lava include clinopyroxene, orthopyroxene, and amphibole. Plagioclase phenocrysts are absent, and large plagioclase (andesine) and quartz xenocrysts are abundant. The disequilibrium nature of the xenocrysts is shown by sieve textures and labradorite rims on the andesine xenocrysts, and by embayed brown glass rims and clinopyroxene aggregates around quartz xenocrysts. Stop 2-5: Tuzgle Volcano and Middle Miocene Ignimbrite Flows Directions: Continue south on dirt road; stop at 24°07′01″S; 66°27′45″W; 4355 m asl. View to the north of Tuzgle volcano in the Pastos Chicos– Aguas Calientes depression. The depression is bounded by Ordovician magmatic and sedimentary sequences on the east and by the Miocene clastic-pyroclastic facies of the Pastos Chicos Formation and the deformed sediments of the underlying Trinchera Formation on the west. The Tuzgle Young andesite lava flow is well seen from here (see also satellite image in Fig. 2.2). The flow has a blocky to aa structure, pressure ridges developed transverse to the flow direction, and a craggy front. The outcrop to the west is a middle Miocene dacitic ignimbrite flow with a K/Ar age of 15.2 ± 0.5 Ma (Aquater, 1980) that erupted over Ordovician Faja Eruptiva granodioritic porphyries (Fig. 2.4). This moderately welded ignimbrite is crystal rich and has a moderate lithic content. The Miocene volcanic rocks of this age that extend as far east as 65°50′W at this latitude (Hongn and Seggiaro, 2001) are argued by Coira et al. (1993) and Kay et al. (1999) to record the initial steepening of the Miocene flat slab in the northern Puna. Stop 2-6: Abra del Charcos and Structure of the Eastern Cordillera Directions: Continue south; stop at 24°09′01″S; 66°24′44″W; 4350 m asl. Looking to the southeast from this stop, low-grade metasediments of the Precambrian–lower Cambrian Puncoviscana

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Formation and quartzites of the middle-upper Cambrian Meson Group can be seen thrust over synrift facies of the Cretaceous Pirgua Subgroup. Just to the west of this fault, Pirgua Subgroup sandstones and conglomerates can be seen in high-angle fault contact with Ordovician Faja Eruptiva Oriental porphyries. To the east of the fault, Ordovician sediments are thrust over Precambrian and Cambrian sequences in a younger over older westvergent thrust that records the inversion of the normal faults of the Cretaceous rift structures. These relations are shown on the geologic map in Figure 2.4. This sequence of structures records the complex history of Late Ordovician (Ocloyic) west-vergent contractional deformation, Cretaceous-Paleogene rifting, and Miocene-Pliocene contractional deformation in this region. Outcrops along the road at this point are Ordovician Faja Eruptiva Oriental granodioritic porphyries, which are distinctive for their large feldspar megacrysts. Stop 2-7: View of the San Jerónimo and Negro de Chorrillos Shoshonites along the Calama-Olacapato-Toro Fault Zone Directions: Continue south on Route 74. Stop is at 24°12′37″S; 66°24′24″W; 4118 m asl. This is a complex interchange; follow the road to Santa Rosa de los Pastos Grandes (Route 129). The monogenic San Jerónimo and Negro de Chorrillos shoshonitic volcanoes seen to the south and southeast from this stop erupted along northeast-trending, strike-slip faults (Fig. 2.5). K/Ar ages are 0.78 ± 01 Ma for San Geronimo and 0.2 ± 0.08 Ma for Negro de Chorrillos (Aquater, 1980). The northwest-trending Chorrillos fault cutting the northern part of the Negro Chorrillos center and the Incahule fault (Petrinovic, 1999) south of the San Jerónimo center define the Calama–Olacapato–El Toro fault zone (megatraza del Toro) in this part of the plateau. The latest movement on the Chorrillos fault was characterized by oblique normal–left-lateral motion (Marrett et al., 1994). The shoshonitic centers are composed of block and aa-type lavas, bombs, scoria, and ash-fall deposits. Lavas from San Jerónimo flowed up to 10 km from the vent; those from Negro de Chorrillos flowed some 4 km down the valley. Ash-fall deposits up to 10 cm thick from the San Jerónimo center are observed in Quaternary fluvial terraces. Déruelle (1991) suggested the shoshonites formed from melts of mica-bearing peridotites that were substantially contaminated in the crust. Knox et al. (1989) and Coira and Kay (1993) argued that the shoshonitic magmas formed as small-percentage melts of a mantle wedge contaminated by subduction-modified lithosphere followed by some degree of crustal contamination. Kinematic studies on the faults in the Quebrada del Toro by Marrett and Strecker (2000) reveal a complex regional structural history. This history involves northwest-southeast contraction on northeast-trending faults that began in the Miocene and ended after 0.98 Ma, and horizontal northeast-southwest contraction on northwest-trending faults that had initiated by 4.17 Ma and is still

active today. Both fault regimes were active between 4.17 Ma and 0.98 Ma. The present kinematic regime along the northweststriking Calama-Olacapato-Toro volcanic belt is controlled by active northeast-striking strike-slip and northwest-trending thrust faults (Fig. 2.5). The two fault sets accommodate the differential shortening between major north-striking thrust fault systems. Stop 2-8: Tajamar Valley near Eastern Topographic Rim of Aguas Calientes Caldera Directions: Stop along Route 129 at 24°14′30″S; 66°26′54″W; 4228 m asl. Three distinctive volcanic units can be seen at this stop. The first is the lower Ordovician Faja Eruptiva granodiorite porphyry with large feldspar megacrysts that is typical of the basement in this region. The second is the “blocky” and intermediate aa-type shoshonitic lava flows from the San Jerónimo center. The third is the highly welded Tajamar ignimbrite (Coira and Paris, 1981), which has a mean K/Ar age of 10.6 Ma, and is the intracaldera facies related to the Aguas Caliente caldera (Petrinovic et al., 1999) to be seen at the next stop. Stop 2-9: View of Rim of Aguas Calientes Caldera Directions: Continue on Route 129; stop near km 11 at 24°17′16″S; 66°27′07″W; 4473 m asl. From this point, the eastern rim of the Aguas Calientes caldera is visible to the east, and the 10.6 Ma Tajamar ignimbrite is seen to overlie the hydrothermally altered Verde and Aguas Calientes ignimbrites on the caldera rim. The older Aguas Calientes ignimbrite has been dated at 17.5 ± 0.5–16.8 ± 0.5 Ma (Petrinovic et al., 1999). The hydrothermal alteration of the older ignimbrites is attributed to the intrusion of subvolcanic dacitic domes. This alteration is responsible for the epithermal Ag-PbZn mineralization exploited in the La Poma mining district near Stop 2-8 (Fig. 2.4). Petrinovic (1999) argues that the collapse of the Agua Calientes caldera was controlled by left-lateral strike-slip faults related to the northwest-southeast–striking Calama-OlacapatoToro fault system. If so, this fault system must have been active at 11–10 Ma. The initial collapse was asymmetric and produced the distinctive rim that is seen only on the eastern side of the caldera. The continuous collapse and opening of rim vents facilitated the eruption of the dacitic Chorrillos, Tajamar, and Abra de Gallo ignimbrites at 10.8 Ma to 10 Ma (Petrinovic et al., 1999). This caldera collapse was followed by a period of resurgence that led to tumescence of the Tajamar ignimbrite intracaldera deposits. Sb-Au epithermal mineralization in the Incachule mining district is associated with the final stages of the volcanic complex. Stop 2-10: Abra del Gallo Ignimbrite Directions: Stop along road at 24°20′20″S; 66°29′28″W; 4691 m asl in the Abra del Gallo.

Neogene evolution: central Andean Puna Plateau and southern-central volcanic zone View of the Abra del Gallo ignimbrite, which is the extracaldera ignimbrite facies associated with the last stage of collapse of the Agua Calientes caldera (Petrinovic, 1999). The eastsoutheast rim of the caldera and related proximal deposits can be seen from this stop. Ash-flow deposits that are up to 170 m thick comprise the ignimbrite plateau to the east and southeast of the rim. In contrast to the Tajamar ignimbrite (intracaldera facies), the Abra del Gallo ignimbrite (extracaldera facies) seen from here is crystal rich, has a low degrees of welding, and lacks vapor phase or hydrothermal alteration. Stop 2-11: Quevar Stratovolcano and Serranía de Barreal Directions: Stop is along road at 24°27′59″S; 66°39′25″W; 3941 m asl. Quevar stratovolcano is a large late Miocene composite center that overlies Ordovician basement. It is the largest volcanic edifice in the west-northwest–trending Calama-OlacapatoEl Toro volcanic chain. The center has been dated and studied by Goddard et al. (1999). At this stop, the andesitic lava flows from the Azufre center (K/Ar age of 7.53 ± 0.02 Ma), and the ca. 5.5 Ma Cerro Gordo dacitic dome can be seen. The lavas cover part of the extensive 10.3 ± 0.3 Ma rhyodacitic-dacitic Olacapato and Pastos Grandes ignimbrites related to the Aguas Calientes caldera. These ignimbrite sheets can be up to 400 m thick under the northern and southeast flanks of the Quevar center. Volcanic activity at Quevar began with the emplacement of the nonexplosive rhyolitic lava domes to the west of Cerro Azufre at 8.6 ± 0.03 Ma. These domes were then covered in large part by major andesitic-dacitic lava flows and minor block and ash deposits that erupted from five vents between 8.3 and 7.53 Ma. The highest part of Cerro Quevar (6180 m) is a dacitic dome dated at 5.58 ± 0.03 Ma. This dome along with the ca. 5.5 Ma Cerro Gordo dacitic dome mark the last magmatism directly related to the Quevar complex. Important hydrothermal alteration and epithermal deposits occur along northeast- and westnorthwest–trending faults in the complex. The largest alteration zone is on the western flank of the volcano in the Quebrada Incahuasi, where a Pb-Ag mine and a precious metal epithermal deposit prospect are located. Looking eastward from this stop, the Tajamar ignimbrite can be seen to cover deformed continental Eocene-Miocene clastic sedimentary sequences. Because the Tajamar ignimbrite is undeformed, the subhorizontal west-northwest–directed shortening and subvertical extension that affected these sediments must have occurred before 10 Ma (Marrett et al., 1994). Stop 2-12: Pastos Grandes Basin Directions: Stop along road at 24°41′49″S; 66°41′37″W; 3880 m asl. The Pastos Grandes basin is located between two major transverse Puna volcanic chains—the Calama–Olacapato–El Toro chain to the north and the Ratones-Archibarca-Galán chain to

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the south. The excellent exposures of the Paleogene to Neogene sediments and Miocene borate deposits in the Pastos Grandes basin near this stop provide a record of the Cenozoic evolution of the Puna. Two principal sedimentary cycles are recorded (Alonso et al., 1991; Alonso, 1999a). The first cycle coincides with deposition of the Paleogene Gestes Formation (Teg; Fig. 2.4) red beds, which contain a subtropical continental fauna and indicate an exoreic fluvial environment with warm and humid climatic condition. These sediments accumulated in broken foreland basins. The second cycle coincides with the deposition of the Pozuelos and Sijes Formations (Tmp and Tms; Fig. 2.4), which formed in endorheic basins (closed with no outflow to rivers) under semiarid to arid conditions. The basins have evaporitic and alluvial deposits (wet salar) along with significant volcanic and geothermal deposits. These early Miocene to late Pliocene (or early Pleistocene) sediments were deposited in compressional retroarc basins in which volcanic centers created barriers (Jordan and Alonso, 1987; Kraemer et al., 1999; Voss, 2002). The presence of Oligocene?-Miocene evaporites shows that the onset of hyperarid conditions was associated with establishment of internal drainage (Alonso et al., 1991; Vandervoort et al., 1995). Strata representing these sedimentary cycles are well exposed near Stop 2-12. The Eocene Geste formation (Turner, 1961; Teg in Fig. 2.4) is a more than 2-km-thick sequence of alluvial and fluvial deposits. These deposits can be seen on the eastern side of the Salar Pozuelos to the west of this stop where they are in angular unconformity with lower Paleozoic metamorphic basement. The exposure is an eastward-dipping homoclinal section of mammal fossil-bearing red and purple conglomerates, sandstones, siltstones, and mudstones. The Geste sequence gives way to the east to the strongly folded clastic and evaporitic (halite, gypsum, and borate) sedimentary beds of the Pozuelos Formation (Turner, 1961; Tmp in Fig. 2.4). These strata are well exposed in the homoclinal structure along the southeastern margin of the Salar de Pastos Grandes. They have yielded a K/Ar age of 7.6 ± 1.1 Ma (Alonso et al., 1991). Above and to the east of the Pozuelos Formation is a thick sequence of siltstones, clay stones, and tuffs in the Sijes Formation (Turner, 1961; Tms in Fig. 2.4). The tuffs have K/Ar ages of 6.8–4.0 Ma (Alonso et al., 1991). North- and northnortheast–trending folds in the Sijes Formation show that active shortening was still occurring at 4 Ma. The Sijes formation hosts an approximately north-northeast–oriented belt of discontinuous borate deposits that are mined in hydroboracite deposits in the Monte Amarillo and Monte Azul mines and in colemanite deposits at Esperanza, Sol de Mañana, and Santa Rosa (see Fig. 2.8). Concordantly above the Sijes Formation is a conglomeratic and psammitic sequence known as the Siguel Formation (Alonso and Gutierrez, 1986; Tps on Fig. 2.4). An intercalated tuff in this sequence yielded a 40Ar/39Ar age of 2.9 ± 0.04 Ma (Vandervoort et al., 1995). Deformation in the Siguel Formation records a period of tectonic reactivation. Terraces made of clastic and evaporitic deposits interbedded with pyroclastic rocks record the presence of a Pleistocene salar

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within and on the southern margin of the Salar de Pastos Grandes (Blanca Lila Formation; Alonso and Menegatti, 1990). The terraces have been dated at 1.5 Ma (see Alonso et al., 2006). Stop 2-13: The Salar de Ratones Depression and the Ratones Volcano Directions: 25°02′42″S; 66°47′5″W; 3870 m asl. Cerro Ratones, which is prominently seen to the south from this stop, is a partially eroded andesitic stratovolcano. This volcanic center is part of the transversal Puna volcanic chain that includes the Archibarca volcano and the Cerro Galán ignimbrite complex (e.g., Alonso et al., 1984b). A K/Ar age of 30 ± 3 Ma (Linares and González, 1990) on an andesitic flow on Cerro Ratones suggests the center is one of the oldest Cenozoic volcanoes in the region. The 40Ar/39Ar age of 7 Ma age in Vandervoort et al. (1995) on a younger flow indicates a later stage of activity. Basement orthogneises and subordinate amphibolites and migmatites on the eastern slope of the Salar de Ratones and in the Salar de Diablillos have yielded a Sm-Nd mineral isochron age of 509 Ma (Lucassen et al., 2000) and a U-Pb age of 508 Ma (Becchio et al., 1999). A Silurian thermal overprint near 440 Ma has been reported by Lucassen et al., (2000). To the west, the metamorphic basement along the western margin of the Salar de Ratones depression consists of quartzaluminosilicate–bearing schist and granitic orthogneiss with subordinate pegmatite, amphibolite, and migmatite (Viramonte et al., 1976). An Ordovician thrust puts these high-grade units over Ordovician sedimentary rocks (Hongn, 1994; Mon and Hongn, 1991). These Paleozoic rocks are unconformably overlain by the Eocene Geste Formation, which is in turn under a tilted and deformed middle Miocene clastic and pyroclastic sequence with north-northeast–trending fold axes. The Miocene deformation is consistent with subhorizontal WNW-ESE shortening and subvertical extension (Marrett et al., 1994). Stop 2-14: Salar del Hombre Muerto Depression (View to the South) Directions: Stop is along road at 25°11′17″S; 66°58′04″W; 4041 m asl. The Salar del Hombre Muerto contains borate deposits interbedded in its sedimentary fill. Borate-bearing salars like these are principally along the eastern border of the southern Puna and in the northern Puna. Ulexite, which is the dominant boron-bearing mineral, occurs in “potatoes” and “bars.” In some salars, borax (or “tincal”) occurs as disseminated (Diablillos and Centenario) or small (Turi Lari, Cauchari, and Rincon) deposits. The formation of Tertiary and Quaternary borate deposits in the Puna can be attributed to three factors (Alonso, 1999a, 1999b): (1) active hot springs with borate-bearing water, (2) closed basins, and (3) an arid to semiarid climate. Borate deposits in the Salar del Hombre Muerto are concentrated on the island in the salar called the Farallon Catal,

and on the Tincalayu and Hombre Muerto peninsulas. The Farallon Catal has a 5000-m-thick sedimentary sequence that is mapped in the Catal Formation (Tmc on map in Fig. 2.4). This formation is composed of red sandstones and pelites at the base, green pelites with gypsum and travertine in the middle, and ignimbrites and reworked tuffs and coarse conglomerates at the top. K/Ar dates limit the age at the base to be 15.0 ± 2.4 Ma and that at the top to be 7.2 ± 1.4 Ma (Alonso et al., 1991). Evaporites near the middle of the section are pure gypsum interbedded with gypsiferous clay (Alonso and Gutierrez, 1986). A basaltic andesite flow overlying the sequence has a 40Ar/39Ar age of 0.8 ± 0.1 Ma (Risse et al., 2008). The folded sedimentary sequence on the Tincalayu peninsula in the northwestern part of the salar is mapped in the Sijes Formation (Tms) in Figure 2.4. This sequence has halite near the base, borax and gypsum in the middle, and clastic strata at the top. A tuff in the borax deposits has an age of 5.86 ± 0.14 Ma. Basaltic andesitic lavas overlying these strata have ages of 2.43 ± 0.7 Ma (Risse et al., 2008) and 0.75 ± 0.3 Ma (Alonso et al., 1984b). The Tincalayu mine contains the largest borax deposit in the Southern Hemisphere. Stop 2-15: Southern Puna Mafic Andesitic Flow Southwest of the Salar de Hombre Muerto Directions: Stop is along side of dirt road that is route 43 near 25°27′30″S; 67°11′41″W; 4115 m asl. The southern Puna is characterized by late Neogene strikeslip and extensional faults, small discontinuous diachronous basins, and mafic, glassy lava flows. Basaltic andesitic to andesitic flows around the Salar de Hombre Muerto like the one at this stop erupted from monogenetic to polygenetic centers along north-northeast–, north- and northeast-oriented faults (Figs. 2.6 and 3.1). They have ages of near 4.6, 2.4, and 0.8 Ma (Risse et al., 2008). The mafic andesite lava (~54% SiO2) at this stop is typical in having high MgO, Cr (300 ppm), and Ni (100 ppm) contents and containing acidic plagioclase and quartz xenocrysts. These features fit with the rapid rise of mantle-derived basaltic magmas (~80%) that mixed with lower crustal melts (~20%) accumulated at mid-crustal depths (Kay et al., 1999, 2005). The quartz and feldspar xenocrysts can be explained as the phenocrysts in the silicic magmas. Mid-crustal mixing of this type is supported by seismic anomalies interpreted to be magma accumulations at mid-crustal depths (e.g., Chmielowski et al., 1999; ANCORP, 2003; Fig. O.5). DAYS 3 TO 5—SOUTHERN PUNA The objective of the next three days is to examine characteristic features of the southern Puna and the differences with the northern Puna. On the afternoon of Day 5, the route descends the southeastern margin of the plateau and crosses into the northern Sierras Pampeanas. A four-wheel drive vehicle is required for many of the stops on Days 3–5.

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Figure 2.6. Thematic Mapper image of a portion of the region shown on the geologic map in Figure 2.4. The topographic margin of the Aguas Calientes caldera is outlined. Quevar is the major late Miocene stratovolcanic complex seen at Stop 10.

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Figure 2.7. Thematic Mapper image of the portion of the map in Figure 2.4 between Volcan Quevar and the Salar de Hombre Muerto. The locations of Stops 2-12 to 2-15 are shown.

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3 km salars, volcanics alluvium Pleistocene Blanca Lila Fm. undif. Tertiary sediments

volcanics Pucara, Rumibola Fms., upper Tertiary Singuel Fm. Pliocene Sijes Fm. Miocene

Fm. Pozuelos (evaporites and clastics) Miocene Geste. Fm. lower member Geste. Fm. middle member Geste. Fm. upper member

Ordovician Fm. Oire Ordovician Fm. Copalayo

Figure 2.8. Geologic map of Santa Rosa de Pastos Grandes region showing the principal borax mines and sedimentary units of the region (Alonso, 1999a).

DAY 3—OLIGOCENE TO QUATERNARY EVOLUTION OF THE SOUTHERN PUNA The objective of Day 3 is to examine the characteristic structural and magmatic features of the southern Puna and the differences in the Miocene to Recent history with the northern Puna (see Overview). A main structural difference is the normal and strike-slip faults along which mafic magmas erupted. Another difference is that the distribution of volcanic centers contributed

to a distinctive sedimentation pattern in a broken foreland basin. Magmatic differences with the northern Puna include (Figs. O.6–O.8): (1) lack of a Neogene magmatic gap, (2) large 15–8 Ma andesitic to dacitic backarc stratovolcanoes and dome complexes, (3) post–7 Ma mafic lavas and glassy andesites and dacites along faults, (4) small- to moderate-sized ca. 6.6–4 Ma ignimbrites followed by the 1000 km3 Cerro Galán eruption at ca. 2.2 Ma, and (5) young mafic lavas with intraplate chemical affinities. These differences are interpreted to reflect moderate

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Figure 2.9. Top: Cross sections through the Tincalayu Mine on the northern margin of the Salar del Hombre Muerto and through the Pastos Grandes basin (Alonso, 1999b). Bottom: Regional east-west cross section at 25°20′S from the Salar de Antofalla to the Salar de Hombre Muerto in the Puna and continuing eastward across the Eastern Cordillera. Profile is from the 1:250,000 Cachi geologic map sheet of Hongn and Seggiaro (2001).

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shallowing of the subducting slab as the Juan Fernandez ridge subducted beneath the region at ca. 14–7 Ma, steepening of the slab after 7 Ma, and delamination of eclogitic lower crust and mantle lithosphere. Limited crustal shortening east of the plateau is consistent with crustal flow contributing to crustal thickening of the southern Puna during the ca. 10–4 Ma peak of ignimbrite eruption in the northern Puna and shallowing of the Chilean flat slab to the south.

Stop 3-1: Outflow Sheet of the Cerro Galán Ignimbrite ~30 km West of the Caldera Rim Directions: Take Route 43 (gravel road) north from Antofagasta de la Sierra toward Salar de Hombre Muerto. Stop is along Route 43 at 25°54′58″S; 67°21′49″W; 3681 m asl. The outflow facies of the large-volume dacitic Cerro Galán ignimbrite that erupted at ca. 2 Ma are well seen at this stop.

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Figure 3.1. Thematic Mapper image of the southern Puna from the Salar de Hombre Muerto to south of the Salar de Antofalla showing Stop 2-15 for Day 2 and all of the stops for Day 3. Important features on the image are the Salars de Hombre Muerto and Antofalla, the various Miocene volcanic complexes, the young mafic flows, and the western part of the Cerro Galán ignimbrite complex and caldera.

Neogene evolution: central Andean Puna Plateau and southern-central volcanic zone The flow at this locality is moderately welded and exhibits spectacular columnar jointing that reflects the geometry of the paleophreatic surface. The ignimbrite overlies a strongly folded Ordovician sedimentary sequence.

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To the east-southeast, the dome structure of the resurgent center in the caldera is seen to rise to an altitude near 6300 m. The Cumbres de Luracatao peaks seen to the north (see Thematic Mapper image in Fig. 2.6) expose the lower Paleozoic orthogneisses and granitoids basement that is under the Cerro Galán caldera.

Stop 3-2: View of Cerro Galán Caldera from the Northwest Directions: Stop is along Route 43 at 25°46′38″S; 67°16′22″W; 4310 m asl. This stop affords a panoramic view of the Cerro Galán ignimbrite sheet that erupted in a major explosive episode at ca. 2 Ma from the giant Cerro Galán caldera (35 × 20 km in diameter) to the east. The ignimbrite has an estimated volume of 1000 km3 (Sparks et al., 1985). The volume of the outflow facies is estimated at 280 km3. The outflow sheet is from 30 to 200 m thick and extends up to 100 km from the caldera rim in all directions. The distal flows visible at this stop cover folded Ordovician sedimentary rocks. Looking to the south, early Pliocene dacitic to rhyodacitic lavas from volcanic centers like Merihuaca (4.86 Ma) can be seen to be partially covered by the Cerro Galán flows.

Stops 3-3a and 3-3b: Views of Beltran Stratovolcanic Complex Directions: Stops are along Route 43 at (a) 25°38′56″S; 67°13′58″W; 4351 m asl; and (b) 25°34′41″S; 67°13′50″W; 4351 m asl. The Beltran stratovolcano seen from Stop 3-3 is one of the large, long-lived composite volcanoes in the southern Puna. The center is dominantly composed of voluminous dacitic to andesitic lava flows and was active from 14.1 to 7.7 Ma (Kraemer et al., 1999). This center, like the Tebenquicho and Antofalla stratovolcanic complexes in the distance to the northwest and west on the margin of the Salar de Antofalla margin (Figs. 3.1 to 3.3), are part of the northwesttrending transcurrent Archibarca–Galán Puna volcanic chain.

Figure 3.2. Portion of the Servicio Geológico Minero Argentino (SEGEMAR) Cachi geologic map (Hongn and Seggiaro, 2001) originally mapped at a scale of 1:250,000 with the locations of Stops 3-1 to 3-4 and 3-7 and 3-8 (black on white background). Stops 3-5 and 3-6 are just to the west of the map and are shown on Figures 3.3 and 3.4.

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Figure 3.3. Generalized geologic map for the central part of the southern Puna from B. Coira showing the region around the Salar de Antofalla. Major salars, volcanic centers, and ignimbrites are identified. Compare with Thematic Mapper satellite image in Figure 3.1. Locations of Stops 3-4 to 3-8 on Day 3 as well as Stops 4-5 to 4-9 on Day 4 and Stops 5-1 to 5-4 on Day 5.

Stop 3-4: Laguna Caro Depression and Glassy Late Miocene Andesites Directions: Turn west from Route 43 onto a dirt track heading toward Laguna Caro. Stop on the south side of Laguna Cara at 25°34′42″S; 67°17′36″W; 4041 m asl. The Beltran seen at the last stop flanks the southern side of the Laguna Caro depression. The eroded glassy andesitic flows visible at this stop and seen elsewhere around the depression are a distinctive volcanic feature of the southern Puna. These flows are characterized by a glassy character, plagioclase (oligoclase) and quartz xenocrysts, and high Mg numbers and Cr (~100–200 ppm) and Ni (~25–100 ppm) contents for their SiO2 (59%–64%) contents. Moderately steep REE patterns (La/Yb

>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.

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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

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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).

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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

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n

t

o

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LEGEND Lagoon Salars C° Blanco Caldera Complex: <0.2 Ma 0.5-0.2 Ma

Lg. Amarga Ig.

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C° Peinado

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Lg. Purulla

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Cerro Blanco Lava-dome Cerro Blanco Ignimbrite Lg.Amarga-Lg.Verde-Vallecito Ign. Complex Mafic lavas Galán Caldera Complex:

Caldera C° Blanco Lg. Verde

26°30´S

2 Ma 2.1 Ma 2.2 Ma 6.4-4.8 Ma 6.3 Ma 8-7.5 Ma

Postcollapse lava-domes Resurgent Center Ignimbrite

Galán Ignimbrite Merihuaca - Real Grande Ignimbrites Rosada Ignimbrite Los Colorados Ignimbrite Calderas structural limit Lineament or regional fault

0

20 km

Figure 4.1. Map of principal latest Miocene to Quaternary ignimbrites and mafic flows of the southern Puna. The principal centers are the giant late Miocene to Quaternary Cerro Galán caldera complex (Stops 4-1 to 4-4), the Pliocene La Amarga and Vallecito ignimbrite centers (the contemporaneous Laguna Verde ignimbrite in Chile will be seen on Day 6), and the Quaternary Cerro Blanco Caldera complex (Stops 4-6 to 4-9). Map compilation by B. Coira. Ig.—ignimbrite.

Neogene evolution: central Andean Puna Plateau and southern-central volcanic zone This stop provides a closeup view of the wavy upper surface of the ca. 2 Ma Cerro Galán ignimbrite. This ignimbrite is the most extensive and youngest flow unit from the Cerro Galán caldera complex with flows reaching distances of up to 100 km from the rim. At this point we are ~18.5 km west of the rim (see Fig. 4.2). The ignimbrite here is dacitic in composition (~68% SiO2), crystal rich (plagioclase, sanidine, quartz, and biotite), has small amounts of pumice (1–5 cm in diameter), and contains sparse lithic fragments. Stop 4-3: Real Grande Valley: Cerro Galán Ignimbrite Sequence near Camp 1 Locality of Sparks et al. (1985) Directions: Continue on dirt track. Stop is near end of track at 25°59′23.6″S; 67°13′57.4″W; 4212 m asl. View is that in Fig. 4.4A. A complete section of the Cerro Galán ignimbritic sequence is exposed in the Real Grande Valley at this stop. A labeled photo looking to the north is shown in Figure 4.4A. The locality is some 12 km west of the caldera rim and is near section CG2 studied by Sparks et al. (1985; Fig. 4.5). The base of the profile is composed of the Blanco ignimbrite, which is overlain by the 6.4–5.1 Ma Merihuaca ignimbrite. The Merihuaca ignimbrite is separated

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into lower, middle, and upper members at this point. The next unit is the 5.14–4.8 Ma Real Grande ignimbrite. The Merihuaca and Real Grande ignimbrites units are characterized by basal plinian deposits, numerous individual flows, and proximal co-ignimbrite lag deposits. Overall, they are moderate to rich in pumice and have a high lithic and a moderate to low crystal content (plagioclase, quartz, biotite, and absent to minor hornblende). Variable compositions are indicated by banded pumices that occur at the top of some units. A dense rock equivalent (DRE) volume of >500 km3 was estimated for the Merihuaca and Real Grande ignimbrites by Sparks et al. (1985). At the top of the section is the ca. 2 Ma Cerro Galán ignimbrite that is separated from the Real Grande ignimbrite by an important erosional unconformity. As at previous stops, the Galán ignimbrite is a crystal-rich, pumicepoor, and lithic-poor unit. In contrast to the underlying ignimbrites, the Cerro Galán ignimbrite lacks basal plinian deposits, proximal lag breccias, and compositional zoning. Sparks et al. (1985) proposed that the Cerro Galán ignimbrite eruption resulted from a catastrophic foundering of a cauldron block into a magma chamber leading to caldera collapse. Post-caldera eruptions took place along the northern segment of the fault on the eastern side of the caldera and on the western margin. Resurgent doming centered along the eastern fault is shown by radial tilting of the ignimbrite and lake deposit accumulations. Stop 4-4: Merihuaca Ignimbrite Section in the Real Grande Valley

67° W

Directions: From near Stop 4-3, head north-northeast on a minor dirt track and continue cross country toward the section viewed at Stop 4-3. Do not do this without a good four-wheel drive vehicle and an experienced driver. Stop is at 25°58′46.7″S; 67°13′08.6″W; 4305 m asl. After the stop, retrace the trip route to the village of Antofagasta de la Sierra.

Volcan Beltran

Merihuaca

4-1

4-2

Antofagasta de la Sierra

4-4 Real Grande 4-3

26°S Cerro Galan Caldera

5 km mafic lavas

Merihuaca

north

4 3 Camp 1

1 2

26°S

10 km 67°15′ W

Figure 4.2. Thematic Mapper image showing the locations of Stops 4-1 to 4-4 relative to the Cerro Galán caldera, the Miocene Cerro Beltran volcanic complex, and the latest Miocene to Quaternary mafic volcanic flows to the west and southwest.

Figure 4.3. Thematic Mapper image of the western side of the Cerro Galán complex providing a closeup view of the region of Stops 4-1 to 4-4. The Laguna Diamonte is well seen on the east side of the caldera wall. Stops 3 and 4 are near Camp 1 of Sparks et al. (1985).

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Figure 4.4. (A) Photograph shows the Cerro Galán ignimbrite sequence in the Quebrada Real Grande as viewed from Stop 4-3. This is one of the most complete sections of the Cerro Galán ignimbrite and corresponds to the sequence in section CG2 in Figure 4.5 studied by Sparks et al. (1985) and Francis et al. (1989). (B) View looking south through the caldera from the perspective of a Thematic Mapper image superimposed on a SRTM (Satellite Radar Topography Mission) image. The Laguna Diamante seen is on the right. (C) Photograph shows the view looking north toward the Cerro Galán resurgent center and the interior of the caldera from the edge of the southern rim of the caldera.

This stop provides a closeup view of part of the Cerro Galán ignimbrite sequence in column CG2 in Figure 4.5. The closest outcrop is in the lower member of the Merihuaca member (Ml in Fig. 4.5). This ignimbrite is a rhyodacite (68% SiO2) with 10%–15% crystals (biotite, quartz and plagioclase). The lower layer of the unit is a massive pumice flow with a slight reverse grading in pumice clasts and up to 20% of lithic clasts (lava fragments and metamorphic rocks). The overlying layer shows reverse grading in pumice and internal bedding defined by variations in pumice size. Four types of pumice occur: (1) a moderately compact gray to white crystal-rich type, (2) a white laminar type, (3) a finely vesiculated ochre type, and (4) an almost aphanitic type. The top layer is a massive pumice flow deposit that is moderately welded and has an intercalated, fine ash-fall deposit. A walk to the north provides a view of the upper units. Salar de Incahuasi and Cerro Blanco Caldera The second half of Day 4 includes a transit through the Salar de Incahuasi to the Cerro Blanco volcanic complex to look at the structural relations (Stop 4-5) and late Miocene to Pleistocene volcanic rocks of the southernmost Puna (Stops 4-6 to 4-9). The late Miocene Rosada ignimbrite and the Quaternary volcanic rocks

associated with the Cerro Blanco caldera will be visited. The region south and west of the Cerro Blanco caldera in the Cordillera de Buenaventura will be seen on Day 6. The stops are shown on the geologic maps in Figures 3.3 and 4.6 and on the Thematic Mapper image in Figure 4.7. The stops require a four-wheel drive vehicle and can only be reached when conditions are appropriate for driving on the surface of the salar. The road is a track that is poorly marked and hard to follow in some segments. Stop 4-5: Neo-Paleozoic Structures Reactivated by Tertiary Andean Tectonism Directions: Head west from the village of Antofagasta de la Sierra on a road to the Salar de Incahuasi. Ask for directions in the village. Stop 4-5 is at the north end of the salar at 26°6′35″S; 67°35′10″W; 3540 m asl. View of the Sierra Phyllo Colorado, which exposes red Permian conglomerates and sandstones (Patquia and de la Cuesta Formations; Fernández Seveso et al., 1991) and Ordovician turbidites (Fig. 4.6). A NNW-SSE–striking normal fault in the northern part of the exposure puts the Ordovician sequence in fault contact with the Permian sequence. These Permian deposits show an abrupt decrease in thickness west of the Sierra Phyllo

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

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Toconquis Fm. and Cerro Galán Ignimbrite Stratigraphy Western flank

Eastern flank

Cerro Galán

Cerro Galán

Unconformity

Unconformity

Real Grande

Cueva Negra

Upper Merihuaca

Upper Leon Muerto

Middle Merihuaca

Lower Leon Muerto

Lower Merihauca Blanco

Region of Stops 4-3 & 4-4

after Sparks et al. (1985)

Distance from caldera rim (km) Figure 4.5. Summary of the stratigraphy of the Cerro Galán Complex ignimbrites on the eastern and western flanks of the caldera. The pre–Cerro Galan ignimbrite units with ages from ca. 7 to 4 Ma were assigned to the Toconquis Formation by Sparks et al. (1985). The bottom diagram shows the stratigraphic sequences measured by Sparks et al. (1985) at various distance from the caldera rim. The section seen at Stops 4-3 and 4-4 corresponds to column CG2 at a distance of ~13 km from the caldera rim.

Colorado where they are exposed in an upthrown block of an east-dipping normal fault. Both the Ordovician and Permian units are covered in low-angle unconformity by Paleogene (Eocene) Geste Formation reddish to brown sandstones, conglomerates, and pelites. The Geste Formation was seen in the Pastos Grandes Basin on Day 2. The Paleogene and Permian are also related by a reverse fault. This structural picture has been interpreted to result from Andean contractional inversion of extensional structures formed in the Permian basin (Seggiaro et al., 2002).

Stop 4-6: Rosada Ignimbrite Directions: Stop is at 26°26′39.6″S; 67°41′21.0″W; 3254 m asl. The Rosada ignimbrite (8.1 ± 0.5 Ma, Kay et al., 2006; 6.3 ± 0.2 Ma, Kraemer et al., 1999) seen at this stop flowed from the south through the Incahuasi salar depression and partially covered Ordovician sedimentary sequences on the eastern flank of the Sierra de Calalaste (Figs. 4.6 and 4.7).

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Pz

in

67°45′W Neo Neo

4-5 Neo Q Neo

26°15′S Pz u p

Pz in

Q

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Q

Q Pz in

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IgQ Q

IgR IgR

4-9

Scale

IgQ Qv Qb

Qv

IgPp

Pz in

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4-8 Pz in

Q

Qv

Pz up

4-6

Q

Paleog

Pz in

Cerro Blanco

10 km

Q Q

Q

IgQ Neo

Neo

P

Figure 4.6. Geologic map of the Salar de Incahuasi and Cerro Blanco caldera region from the 1:250,000-scale map of the Paso de San Francisco (Seggiaro et al., 2002) showing Stops 4-5 to 4-9. Compare with Thematic Mapper image in Figure 4.7. Labels are for lower Paleozoic units (Pz in), upper Paleozoic units (Pz up), Paleogene sedimentary rocks (Paleog), Miocene to Pliocene units (Neo), late Miocene Rosada ignimbrite (IgR), Quaternary Blanca (White) ignimbrite (IgB) and Piedra Pumice (IgPp), Quaternary ignimbrite (IgQ), Quaternary mafic lava flows (Qb), and other Quaternary deposits (Q).

Subsequently, the Rosada ignimbrite was eroded and partially covered by the Pleistocene Blanco ignimbrite that flowed through the low part of the center of the depression. The Rosada ignimbrite is best exposed on the border of the Blanco ignimbrite along the margins of the Incahuasi depression. The Rosada ignimbrite is a highly to moderately welded rhyodacitic ash flow (66.8%–68.7% SiO2) that is crystal rich, has moderate to large amounts of pumice (15%–35%), and is poor (<5%) in lithic fragments (dacites and Ordovician rocks). The pumice is white, crystal rich, and variably flattened depending on the degree of welding. Some very flattened pumice fragments are up to 15 cm in length. Stop 4-7: Blanca (White) Ignimbrite Directions: Stop is at 26°31′06.7″S; 67°42′15.4″W; 3466 m asl. The Blanca ignimbrite (White ignimbrite) exposed here has a 40Ar/39Ar age of 0.2 ± 0.1 Ma (Siebel et al., 2001) and is related to the Cerro Blanco caldera complex (Figs. 4.6 and 4.7). The complex consists of two nested calderas embedded in a major structure with a diameter of ~15 km. The main ash flows form two units. One is the Campo de la Piedra Pomez unit that flowed northeastward along the Carachipampa valley, where it has been dated at of 0.55 ± 0.1 Ma (Seggiaro et al., 2002). The other is the Blanca ignimbrite seen here. This ignimbrite flowed to the north for more than 25 km through the Incahuasi depression. The flow has a rhyolitic composition (70%–74.4% SiO2) and is massive, crystal poor (quartz, plagioclase, and biotite), pumice rich, and poorly welded. Different facies of the ignimbrite vary from co-lag proximal to distal pumice-rich deposits. The margins of the Cerro Blanco caldera complex are formed by distributed co-ignimbrite breccias that contain older ignimbrite and lava blocks. Several intracaldera lava domes record an episode of magmatic resurgence at 0.15 Ma (Siebel et al., 2001) postdating the Blanca ignimbrite. Pritchard and Simons (2002) reported a thermal anomaly (~100 °C) in the central part of the Cerro Blanco caldera, and detected a subsidence of 2.5 cm for a period from 1992 to 2000 using radar interferometry. Their interferogram is shown as an inset in Figure 4.7. Evidence for subsidence has been absent since 2000. Stop 4-8: Cueros de Purilla Obsidian Lava Dome Directions: Stop at 26°34′3.6″; 67°44′58.2″; 3895 m asl. The Cueros de Purulla is a rhyolitic (72.3% SiO2) obsidian lava dome (Fig. 4.7) with a 40Ar/39Ar date of 0.4 ± 0.1 Ma (Siebel et al., 2001). The location of the dome is controlled by extensional and strike-slip faults. Dome formation can be attributed to the release of batches of degasified magmas at the time of the eruption of the Blanca ignimbrite and the formation of the Cerro Blanco caldera.

10 km Antofagasta de la Sierra

4-5

north

Laguna 5-1 Alumbrera 5-2 5-3

Jote

4

4-6

Cueros de 28°20′ S Purulla

Carachipampa

5-4

4-7 4-8

4-9

Campo de la Piedra Pumice

Cerro Blanco Caldera

Cerro Blanco

67°45′ W

10 km

0 cm 5

Figure 4.7. Thematic Mapper image showing the Salar de Incahuasi, the Cerro Blanco caldera, and the young mafic lavas south of the village of Antofagasta de la Sierra. The western half of the image overlaps the region shown on the geologic map in Figure 4.5. The localities of Stops 4-5 to 4-9 as well as Day 5 Stops 5-1 to 5-5 are shown. See also Figure 3.3. Inset at bottom right shows satellite radar interferogram for Cerro Blanco center from Pritchard and Simons (2002).

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Stop 4-9: Domes at the Margin of the Cerro Blanco Caldera Directions: Stop at 26°40′37.8″; 67°45′6.0″; 4020 m asl. Dacitic lava domes at this stop on the northern margin of the Cerro Blanco caldera complex erupted in episodes of degasification of the caldera system. They are phenocryst-poor vitrophyric lavas (plagioclase, biotite, and quartz). One has a K/Ar age of 1.3 ± 0.4 Ma (Kay et al., 2006) and is considered to constitute the easternmost dome in the east-west–trending Cordon de Buenaventura dome complex. DAY 5—SOUTHERN PUNA YOUNG MAFIC LAVAS AND FAULTS AND THE TRANSITION INTO THE NORTHERN SIERRAS PAMPEANAS Stop 5-1: Quaternary Basaltic Andesite Flows of the Laguna Volcano Directions: Head south from Antofagasta de la Sierra on Route 43 (being improved and paved in late 2007). Stop at 26°07′55″S; 67°25′26″W; 3368 m asl. The Laguna volcano is a prominent monogenetic cinder cone (Figs. 4.7 and 5.1) with a lava flow dated by 40Ar/39Ar at 0.34 ± 0.06 Ma (Risse et al., 2008). The flows have an aa morphology with distinctive jagged flow fronts. In some places, layers of fragmented clinkers alternate with massive lavas. The lavas are characterized by olivine and plagioclase phenocrysts and a basaltic andesitic composition (53% SiO2). Their trace elements show an intraplate-like signature (La/Ta ~25; Ba/Ta ~386) and a moderately steep REE pattern (La/Yb = 17) that is interpreted to reflect an increase in mantle wedge volume above the subducting Nazca plate in response to lithospheric delamination and steepening of the slab (Kay et al., 2005; Overview). Stop 5-2: Late Miocene Glassy Lavas under the Laguna Flows Directions: Stop along Route 43 at 26°08′23″S; 67°25′14″W; 3343 m asl. The glassy andesitic lava flow (~62% SiO2) at this stop, which has a 40Ar/39Ar age of 7.34 ± 0.04 Ma (Risse et al., 2008) is partially covered by the Quaternary Laguna lava flow. The flow shows the characteristic flaggy parting typical of many latest Miocene and Pliocene southern Puna lavas The trace element chemistry contrasts sharply with that of the Laguna flow in having an arc-like signature (strong HFSE element depletion; La/Ta = 46.9). This signature along with a steep REE pattern (La/Yb = 28) showing equilibration with garnet typifies many southern Puna lavas of this age. These characteristics are interpreted to reflect a thick underlying crust and the introduction of arc components into the magma source in association with late Miocene shallowing of the subducting slab under the southern Puna region (Kay et al., 2005).

Stop 5-3: View of the Mafic Jote and Alumbrera Volcanoes Directions: Stop along Route 43 at 26°12′0″S; 67°24′14″W; 3300 m asl. The Jote volcanoes are a cluster of monogenic basaltic andesitic centers composed of cinder cones and lava flows. Their distribution is well seen in the satellite image in Figure 4.7. The centers are distributed along a north-northeast–striking fault system that has a component of horizontal extension. An olivine phyric basalt (51.5% SiO2) from one of these flows yielded a 40Ar/39Ar age of 3.2 ± 0.02 Ma (Risse et al., 2008). The flows are chemically similar to the Laguna flows (Kay et al., 1994a) indicating that generally similar magmatic conditions have persisted in the region for the past 3 million years. Stop 5-4: View of Carachipampa Volcano, Campo de Piedra Pumice, and Cerro Blanco Caldera Directions: Stop is along Route 43 at 26°28′49″S; 67°20′15″W; 3160 m asl. The prominent Carachipampa monogenetic mafic volcanic center is visible to the west from this stop. The Campo de Piedra Pomez, which is the pumice-rich flow erupted from the Cerro Blanco caldera at 0.55 Ma (Seggiaro and Hongn, 1999), is visible to the southwest (Figs. 4.7 and 5.1). The Carachipampa volcanic center is composed of mafic scoria cones and radiating lava flows. The lavas are clinopyroxene-bearing basalts (52% SiO2) with an intraplate-like chemistry (Kay et al., 1999, 2006) that yielded a 40 Ar/39Ar age of 0.75 ± 0.08 Ma (Risse et al., 2008). Stop 5-5: Pasto Ventura Monogenic Volcanoes and Quaternary Faults Directions: Entrance to Pasto Ventura region is along old Route 43 at 26°41′56″S; 67°10′40″W; 3792 m asl. A four-wheel drive vehicle is required. Follow track to south and then turn at 26°44′03″S; 67°11′03″W; 3623 m and drive up gentle slope to west. Stop is at 26°44′24″S; 67°13′08″W; 3675 m asl. As shown in Figures 5.2 and 5.3, this is another region where young southern Puna mafic lava flows are clearly related to normal and strike-slip faults. The locality has been studied by Allmendinger et al. (1989) and Marrett et al. (1994). In the region, Paleozoic metamorphic basement rocks can be seen to be thrust over tightly folded continental clastic sequences that are overturned beneath west-northwest to northwest-dipping and north-northeast to northeast- striking thrust faults. These thrust faults are considered to be Miocene in age. Along strike, these faults and possible new faults offset young basaltic andesitic cinder cones and lavas and alluvial surfaces like those seen at this stop. Kinematic indicators along the reactivated fault provide evidence for oblique normal-right lateral motion with moderately northeast and east-northeast–plunging shortening and subhorizontal north-northwest to north-south extension (Allmendinger et al., 1989). A small eroded basaltic andesitic

Figure 5.1. Regional geologic map of the southern Puna and eastern border with the locations of Stops 5-1 to 5-8. The locations of Stops 5-1 to 5-4 are shown on the Thematic Mapper image in Figure 4.7. Map is from the Servico Geológico Minero Argentino (SERGEMAR) 1:500,000-scale map of the province of Catamarca (Martínez, 1995). Symbols for generalized units are Early Paleozoic schists (Sch-Epa), slate-schist (Slate-Sch Epz), and orthogneiss (Orgneis Epz), Ordovician granitoids (GrOrd), Ordovician sedimentary rocks (Ord), Tertiary clastic sequences (Tc), Miocene volcanic rocks (V-Mioc), Cerro Galan ignimbrites (Ign), Pliocene to Quaternary basalts to andesites (B-A), and Quaternary sediments (Q). Farallon Negro is a Miocene volcanic complex.

10 km

26°40′ S Turn-0ff

North

5-7 5-5 5-6

Figure 5.2. Thematic Mapper image of a portion of the southeastern part and eastern slope of the southern Puna showing the locations of Stops 5-5 to 5-8. Features of the volcanic deposits of the Vicuña Pampa Caldera seen at Stop 5-7 are labeled as is the turn-off for Stops 5-5 and 5-6 in the Pastos Ventura region.

26°50′ S Caldera floor

Vicuña Pampa Caldera

27°00′ S

5-8 67°10′ W

PreCambrian/ Paleozoic basement

Figure 5.3. Map of the Pasto Ventura region from Allmendinger et al. (1989) showing the distribution of young mafic lavas relative to strike-slip and normal faults compared to a Thematic Mapper Image (boxed area shows region of map) showing the locations of Stops 5-5 and 5-6. Age of Quaternary basalt (Qb) is from Marrett et al. (1994).

Miocene sandstone

67°15′ W 1.3±0.6 Ma

north 2 km

5-6

66°50′ W

Miocene volcanics

Qb

5-5

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Neogene evolution: central Andean Puna Plateau and southern-central volcanic zone cone displaced by a normal fault to the north of this stop has a 40 Ar/39Ar age of 1.3 ± 0.6 Ma (Fig. 5.2; Marrett et al., 1994). The young basaltic to mafic andesitic lavas (52%–56% SiO2) in this region contrast with the young mafic lavas erupted at the Laguna, Jote, and Carachipampa centers in having arclike high field strength element depletions (La/Ta = 32–55; Ba/Ta = 430–670; Kay et al., 1994a). Stop 5-6: Late Miocene Stratigraphy and Mio-Pliocene Deformation near Pasto Ventura Directions: Continue south along rough track to 26°45′23″S; 67°13′29″W; 3684 m asl. After stop, retrace route back to main road (Route 43). The late Miocene sequence at this point consists of a red bed sequence containing an ash-fall tuff dated by 40Ar/39Ar at 10.04 ± 0.05 Ma (Marrett et al., 1994). The sequence is affected by overturned folds that are cut by west-northwest to northwestdipping thrust faults. The geometry of the structures indicates subhorizontal west-northwest to northwest-trending shortening and subvertical extension during the late Miocene to Pliocene (Marrett et al., 1994).

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ment rocks over folded Tertiary sequences. Northwest horizontal shortening and vertical extension directions are prevalent (Allmendinger, 1986). En echelon faults terminate in plunging folds like those that are typical in transfer zones of foreland thin-skinned thrust belts. The fault segments in the Sierra Pampeanas are shorter than those fault segments in the Puna (average 20–30 km in length). The timing of contractional deformation is constrained by ages in the Corral Quemado sequence (Fig. 5.4) that indicate that contractional deformation began after ca. 11 Ma and continued until ca. 3 Ma (Allmendinger et al., 1989). The second and younger style of late Cenozoic deformation in the northern Sierras Pampeanas is latest Pliocene-Quaternary in age and is associated with a thrust kinematic regime typified by nearly east-west shortening and vertical extension. This geometry is seen along the Rio El Bolsón where the Punaschotter conglomerates (Fig. 5.4) unconformably overlie folded sequences in the Andalhuala Formation that are cut by northeast-striking thrust faults. Allmendinger et al. (1989) contrast this style with the contemporaneous strike-slip and normal faults on the plateau to the west, which were seen at Stop 5-5 at Pasto Ventura and other southern Puna stops. The southern Puna extensional faults often have a generally northerly strike and dip 0° to 47°N; the strike-slip faults have a nearly east-west shortening direction like the thrust faults in the northern Sierras Pampeanas.

Stop 5-7: View to South of the Vicuña Pampa Volcanic Complex Directions: Stop along Route 43 near 26°44′29.2″S; 67°04′15.7″W; 3897 m asl. The Vicuña Pampa Volcanic Complex seen to the southsoutheast is a succession of ignimbrites, lavas, and minor dikes of basaltic to basaltic andesitic composition associated with rare plinian deposits. The complex is constructed on Precambrian–Lower Paleozoic metamorphic basement. Viramonte and Petrinovic (1999) proposed that this volcanic complex developed as a collapse caldera with a circular geometry and a diameter of 15 km. The elements of the caldera are shown on the Thematic Mapper image in Figure 5.2. An andesitic lava flow associated with the center at Peñas Frias has a K/Ar age of 12.1 ± 0.6 Ma (Kay and Coira, unpublished). A N40°W-striking fault system cuts the complex. A fault displacing the rim is reported to have a left-lateral strike-slip sense with ~3 km of displacement (Allmendinger et al., 1989).

stop 6 Quaternary basalt

tan tuffaceous sandstone

near stop 8 Tertiary-Quaternary conglomerate (Punaschotter) El Cajon Formation (Araucanense tuffaceous sandstone)

redbeds 10.04±0.5

El Morterito Formation (Calchaquense redbeds)

Stop 5-8: Rio El Bolsón—Phases and Styles of Late Cenozoic Deformation Directions: Stop along Route 43 at ~27°01′28″S; 66°45′43″W; 2400 m asl. Two distinct phases and styles of late Cenozoic deformation have been described in the northern Sierras Pampeanas near the boundary with the southern Puna by Allmendinger (1986, 1989). The older phase is characterized by contraction exhibited by folding and dip-slip thrust faults. Northeast-trending thrust faults that dip 20°–55° to the northwest bound structural blocks and put base-

Figure 5.4. Simplified stratigraphic columns and cross sections showing correlations and contrasts in structural style between the Pasto Ventura region in the southern Puna (Stops 5-5 and 5-6) and the Corral Quemado region in the northern Sierras Pampeanas in the Rio Bolsón area (Stop 5-8). Numbers on columns are ages. Figure is modified from Allmendinger et al. (1989). Age on tuff in Pasto Ventura redbeds is from Marrett et al. (1994).

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Farellón Negro

Figure 5.5. Tectonic-stratigraphic summary of the southeastern margin of the Puna from Allmendinger (1986).

At this stop, Andalhuala Formation medium- to coarsegrained sandstones, conglomerates, and siltstones are affected by folds with northeast-trending axes that are typical of the first deformation stage. Elsewhere in the region, northeaststriking faults thrust Paleozoic basement over these units. The Miocene age of the Andalhuala Formation is based on 40Ar/39Ar dates of 6.68–7.14 Ma on interbedded tuffs (Allmendinger, 1986). FIELD GUIDE PART 2: ARC REGION IN CHILE Suzanne Mahlburg Kay and Constantino Mpodozis The objectives of Days 6 and 7 are to examine the Neogene to Recent history of the Miocene to Recent frontal volcanic arc region in Chile. During the late Oligocene to late Miocene (26–6 Ma), the frontal arc was located along the Maricunga Belt and then migrated eastward to the site of the Pleistocene to Recent Central Volcanic Zone arc. The southernmost Central Volcanic Zone Ojos del Salado region is visited on Day 6 and the Maricunga Belt on Day 7. DAY 6—SOUTHERNMOST CENTRAL VOLCANIC ZONE, NEOGENE ARC MIGRATION, AND QUATERNARY TO HOLOCENE ACTIVITY Directions: Take paved Route 45 from Fiambalá through Chaschuil to the Argentine customs and immigration station at Las Grutas just east of Paso San Francisco and the Chilean border. The distance from Fiambalá to Paso San Francisco is ~205 km.

Introduction to the Volcanic Centers of the Ojos del Salado Region The Ojos del Salado volcanic region is located near the southern termination of the modern Central Volcanic Zone near 27°S latitude (Figs. 6.1 and 6.2). This region is home to the Central Volcanic Zone arc and the late Oligocene to Miocene volcanic centers that erupted in the backarc of the Maricunga Belt arc, which is over 40 km to the west (Fig. 6.1). Between ca. 9 and 6 Ma, Maricunga Belt volcanic activity dramatically decreased and then ceased as volcanism increased significantly in the backarc. By 4 Ma, the main frontal volcanic arc was essentially in the Ojos del Salado region. The distribution, age, and geochemistry of the Ojos del Salado region volcanic rocks reflect complex magmatic-tectonic interactions associated with arc migration, crustal thickening, and uplift (Figs. 6.1 and O.10; Mpodozis et al., 1996; Kay et al., 1997, 1999, 2006). Considering the modern subduction geometry and assuming that the magmatic front of the Maricunga Belt was located ~100 km above the subducting Nazca slab, a significant portion of the forearc crust and frontal arc lithosphere should have been removed from below this region after 8 Ma (Kay and Mpodozis, 2000, 2002; Kay et al., 2006). The objectives of Day 6 are to see the Central Volcanic Zone volcanic front, view the Miocene backarc volcanic centers, and discuss the processes associated with eastward arc migration. Most of the volcanic activity in the Ojos de Salado region can be summarized in the three episodes listed below (see Mpodozis et al., 1996). Centers mentioned are labeled on Figures 6.1 to 6.4; K/Ar ages are plotted in Figures 6.5 to 6.7. Discussions of the geology, physical Volcanology, and geochemistry of the centers can be found in González Ferrán et al.

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Figure 6.1. Generalized regional geologic map showing the age distribution of Miocene to Holocene volcanic rocks in the southernmost Central Volcanic Zone (CVZ). Field trip route on Day 6 passes near 27°S latitude. Map is from Kay and Mpodozis (2000).

(1985), Baker et al. (1987), Mpodozis et al. (1996), Gardeweg et al. (1997, 1998, 2000), Kay and Mpodozis (2000, 2002), Kay et al. (1999, 2006), and Clavero et al. (2000). Late Miocene (9–5 Ma) Volcanism behind the Dying Maricunga Belt Arc An older group includes andesites and dacites erupted from isolated 9–7 Ma stratovolcanoes. They include a lava from the

Laguna Verde volcano (8.9 ± 0.4 Ma), flow from the lower part of the Cerro Los Patos volcano in Argentina (7.6 ± 0.6 Ma), and 7.8–7.6 Ma lavas in the Cordon Foerster. A younger group includes an andesite along the Robertson Escarpment (6.5 ± 0.7 Ma), 5.9–5.1 Ma high-K andesites and dacites around the Wheelwright caldera, flow banded rhyolites associated with the Fuenzalida dome (5.8 ± 0.9 Ma), and dacites associated with the La Barda Complex (4.9 ± 0.4 Ma).

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Figure 6.2. Part of the Chilean Geologic Survey (SERNAGEOMIN) geologic map of Chile mapped at a scale of 1:1,000,000 showing the Maricunga Belt and the southernmost Central Volcanic Zone. Major volcanic centers are indicated by name and color coded by age to show the pattern of arc migration. The field trip stops are designed to examine volcanic centers and features related to each of these time periods. The map is available for purchase through the SERNAGEOMIN in Santiago.

Pliocene (5–1.8 Ma) Volcanism at the Emerging Central Volcanic Zone Front After volcanism ended in the Maricunga Belt, the front migrated eastward, and large-scale activity began in the Ojos del Salado region. From 4 to 1.4 Ma, diverse centers erupted moderate volumes of basaltic andesite to rhyolitic lavas. Older units include silicic andesitic flows at Los Patos volcano (4.5 ± 0.9 Ma), silicic andesitic and dacitic lavas and domes southwest of Tres Cruces (4.0–4.5 Ma; Gardeweg et al., 1997), glassy andesitic lavas from the Rodrigo cone (4.4 Ma), and andesites from the Peñas Blancas (4.2 Ma) and Ermitaño volcanoes (3.6–3.7 Ma) and a center north of the Cerro Laguna Verde volcano (3.9 ± 0.9 Ma). The Laguna Verde rhyodacitic ignimbrite erupted at 3.8–4.4 Ma. Younger eruptions include the thick 3.5–3.4 Ma dacitic flows that are the precursors to the Ojos del Salado complex (Fig. 6.6) and pyroxene andesite flows at Cerro El Muertito (3.2 ± 0.3 Ma) and Volcán del Inca (3.3 ± 0.3 Ma). Similar age silicic rocks erupted around the Nevado Tres Cruces (Fig. 6.7).

Pleistocene to Holocene Centers at the Central Volcanic Zone Front (<1.5 Ma) The most striking centers of the Ojos del Salado region are the Quaternary volcanic edifices that form a northeast- to eastnortheast–trending belt that extends for more than 50 km (Fig. 6.4). Most of these centers are well over 6000 m high. Major andesitic to dacitic stratovolcanic complexes near the border include Incahuasi (6610 m), San Francisco, Cerro Condor, and Cerro Peinado. Centers farther east include El Fraile, El Muerto (6420 m), Ojos del Salado (6880 m), Cerro Solo (6190 m), and Nevado de Tres Cruces (6330 m), which are dominantly dacitic dome complexes and lava domes. Centers with K/Ar and 40Ar/39Ar ages less than 1.5 Ma constitute more than half of the erupted rocks in the Ojos del Salado region. The dominance of andesites and presence of basaltic lavas (Incahuasi and San Francisco) in the Paso San Francisco region and dacitic complexes elsewhere can be related to regional differences with contractional stresses dominating in the west and extensional stresses along faults being more important in the east.

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Figure 6.3. Thematic Mapper image shows the southernmost part of the Central Volcanic Zone in the Paso de San Francisco region and the locations of Stops 6-1 and 6-2. Also shown are the Cerro Blanco caldera and Cueros de Purulla dome seen on Day 4.

Stop 6-1: Las Grutas—View of the Latest Pliocene to Quaternary (Holocene?) San Francisco and Incahausi Volcanoes at the Southern End of the Central Volcanic Zone Directions: Stop along Route 45 in Argentina just west of immigration and customs buildings and east of Paso San Francisco (pass at 4725 m) at 26°54′23″S; 68°23′55″W; 4424 m asl. The centers of the Ojos del Salado region near Paso San Francisco (Incahuasi, San Francisco, and Falso Azufre) are typical Andean stratovolcanoes (typically 58%–61% SiO2) that

dominantly erupted pyroxene and hornblende andesites. The Incahuasi and San Francisco volcanoes viewed from this stop also have monogenetic parasitic cones composed of olivine- and clinopyroxene-bearing basalts (~53% SiO2). Kay et al. (1999) interpreted these mafic lavas as products of mantle-derived magmas that were contaminated by siliceous magmas like those from the nearby Pliocene Laguna Verde ignimbrite (Figs. 6.2 and 6.4). Contamination occurred as the magmas ascended rapidly along deep-seated faults. The east-northeast–trending Cordillera de San Buenaventura, just to the north of Las Grutas (Fig. 6.3) includes an

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Figure 6.4. Thematic Mapper image shows the southernmost part of the active Central Volcanic Zone and the forearc, which was the backarc of the Miocene Maricunga Belt arc. Centers in the forearc largely erupted at the time of frontal arc migration. Locations of Stops 6-1 to 6-8 are shown. Blue regions are areas of snow and ice and mark the highest peaks.

east-northeast–trending chain of latest Pliocene to Pleistocene andesitic to dacitic domes with ages ranging from 2.3 to 0.4 Ma (Kay et al., 2006). The east-northeast–trending faults along which these domes were emplaced reflect complexities related to major structural blocks. Baldwin and Marrett (2004) have argued that the faulting and the domes in the Cordillera de San Buenaventura are related to a releasing bend connecting regional-scale, northnortheast–trending, right-lateral strike-slip faults. Stop 6-2: Late Miocene Mulas Muertas Volcanic Center Directions: Head west through Paso San Francisco crossing the border into Chile; stop along Chilean National Route 31 near 26°54′54″S; 67°07′59″W; 4019 m asl. The Mulas Muertas volcanic complex seen directly to the west (Figs. 6.4 and 6.5) from this stop is the largest of the backarc centers that erupted in the Ojos del Salado region as the late Miocene Maricunga volcanic arc was expiring to the west. The oldest part of the complex is a hornblende-biotitesanidine dacite dome dated at 5.6 ± 0.4 Ma. The main erup-

tion of the center produced two-pyroxene andesites and hornblende-bearing andesites and dacites with K/Ar ages of near 5.1–5 Ma. Other late Miocene backarc volcanic rocks in the region include dacitic lavas dated at 6.5 ± 0.7 Ma that along the Robertson escarpment to the west of the San Francisco volcano and andesitic to dacitic rocks with ages from 5.9 to 5.3 Ma near the Wheelwright volcano (see Fig. 6.5). This stop also affords a nice view of the Incahuasi volcano up the valley to the south-southeast and a view of the El Fraile center directly to the south (see Figs. 6.4 and 6.5). Laguna Verde is the lake to the west. Stop 6-3: Laguna Verde and Pliocene Laguna Verde Ignimbrite Directions: Stop along Chilean Route 3 near 26°53′25″S; 68°29′0″W; 4375 m asl. The Laguna Verde ignimbrite (Fig. 6.4) is a thick and widespread, pumice-rich, dacitic to rhyolitic ignimbrite. The upper levels are formed by a biotite-quartz-sanidine–bearing welded

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68°30′ W Caldera Wheelwright 3.9±0.9 <1 5.9±0.7 3.7±0.6

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Figure 6.5. Geologic map of the greater Ojos del Salado–Tres Cruces–Caldera Wheelwright region from Mpodozis et al. (1996) showing major volcanic centers with their K/Ar ages and the major faults in the region.

rhyolitic tuff (>74% SiO2) with K/Ar ages ranging from 4.4 to 3.8 Ma (Fig. 6.5; Mpodozis et al., 1996). The source is uncertain. Since the Laguna Verde ignimbrite is similar in both composition and age to the Laguna Amarga and Vallecito ignimbrites (Siebel et al., 2001; Kay et al., 2006) just to the northeast (Fig. 6.3), the source has been suggested to be related to the Laguna Amarga caldera described by Seggiaro et al. (2002). The Laguna Verde ignimbrite erupted as the volcanic front was moving to the present site in the southernmost Central Volcanic Zone (Mpodozis et al., 1996; Kay et al., 1997).

Stop 6-4: View of Pliocene to Holocene Ojos del Salado Complex Directions: Stop along Chilean Route 31 near 26°54′55″S; 68°35′10″W; 4466 m asl. The Ojos del Salado peak with a height of 6880 m is well seen through the valley to the south from this stop. The Ojos del Salado complex is a large constructional volcanic center built by numerous superposed small volcanic edifices. The center, which has been mapped and described by Gardeweg et al. (1997, 1998;

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see Fig. 6.6), can be seen to be extensively covered by air fall pumice deposits that originated from an explosive eruption from the Nevado de Tres Cruces to the west. The pumice deposits are thickest on the lower slopes. The lowermost part of the Ojos del Salado complex consists of thick dacitic flows that have yielded K/Ar ages of ca. 3.4–3.5 Ma (Mpodozis et al., 1996). These flows are similar in age to low-silica pyroxene andesite lavas dated at 3.2 ± 0.3 Ma at Cerro El Muertito (Figs. 6.4 and 6.5). The main

part of the Ojos del Salado complex is composed of short, thick, dacitic lava flows and domes clustered in an area 13 km long by 12 km wide. The lower part dates to the early Pleistocene as indicated by a K/Ar age of ca. 1.53 Ma from a long, glassy, two-pyroxene andesite flow in the lower part. The upper part is composed of short, thick, youthful-looking glaciated dacitic lava flows. Those on the western flank form a north-northeast– trending line of small-volume, two-pyroxene, hornblende-biotite

Figure 6.6. Geologic map and ages of the Ojos del Salado complex after Gardeweg et al. (1997, 1998) showing principal peaks, elevations in meters, geologic units, and K/Ar and 40Ar/39Ar ages.

Neogene evolution: central Andean Puna Plateau and southern-central volcanic zone dacite cones. A summit complex is developed over northwardextending flows that have yielded a K/Ar age of 1.08 ± 0.09 Ma. The summit complex consists of two coalesced, east-west– aligned edifices whose diameters are less than 2.5 km. The hornblende-biotite dacites in these edifices yielded a 40Ar/39Ar date of 0.34 ± 0.19 Ma (Gardeweg et al. 1997, 1998). The active crater (1.3 × 0.5 km), which shows signs of fumarolic activity, is located between the two edifices. The peaks seen to the south on the west side of the valley constitute the Cordon Foerster. The main center has a late Miocene K/Ar age of 7.5 ± 0.9 Ma (Fig. 6.5). The center on the east side of the valley is the Mulas Muertas volcano that was viewed from another perspective at the previous stop. The view here is of the hornblende-clinopyroxene–bearing andesites and dacites of the main eruptive phase with K/Ar ages of 5.0 ± 0.3 Ma and 5.1 ± 0.3 Ma. Stop 6-5: Southern Rim of the Wheelwright Caldera Directions: Stop along Chilean Route 31 near 26°54′40″S; 68°37′15″W; 4510 m asl. The outcrop along the road at this point is the Pliocene Laguna Verde ignimbrite. The enigmatic Wheelwright caldera seen to the north of the road at this stop is defined by a ~19-km–wide, subcircular, topographic rim that surrounds a depression. The rim and depression are well seen in the satellite image in Figure 6.4. A problem with interpreting this structure is that no major ignimbrites are associated with the caldera-like depression. The main Wheelwright structure formed between 5.2 and 4.2 Ma (Baker et al., 1987; Clavero et al., 2000) just after the Maricunga arc shut off and as the new arc front was starting to form in the position of the modern southern Central Volcanic Zone. A series of Pliocene stratovolcanoes sit along the rim of the Wheelwright caldera (see Fig. 6.5). These include the large Peñas Blancas volcano (4.2 ± 0.6 Ma) on the southwest side (seen to the northwest from this stop), the Ermitaño volcano (3.7 ± 0.6 Ma and 3.6 ± 0.6 Ma) on the southeast side, and a small unnamed center north (3.9 ± 0.9 Ma) of the Laguna Verde volcano (Fig. 6.5). The late Miocene (5.9 ± 0.9 Ma) Pico Wheelwright center sits just west of the Peñas Blancas center. The Volcán del Inca on the south side of the Peñas Blancas center (viewed immediately to the north from this stop) is an olivine-pyroxene basaltic andesite parasitic cone with a K/Ar age of 3.3 ± 0.3 Ma. Stop 6-6: Early Miocene Segerstrom Basalt, Middle Miocene Cerros Amarillos Volcanic Center, and the Pliocene Rodrigo Volcano Directions: Stop along Chilean Route 31 near 26°56′50″S; 68°45′58″W; 4500 m asl. Geological features at this stop include the southwestern rim of the Wheelwright caldera to the northeast (discussed at Stop 6-5), Oligocene to Miocene sedimentary and mafic volcanic sequences of the Cordon Segerstrom to the north, altered middle

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Miocene Cerros Amarillos lavas to the east-northeast, and the early Pliocene Rodrigo Volcano to the south (Figs. 6.4 and 6.5). The view to the north into the Cordón Segerstrom shows the basaltic and basaltic andesitic lavas called the Segerstrom basalt (González-Ferrán et al., 1985; Baker et al., 1987). These lavas cover the late Oligocene–early Miocene volcanic and sedimentary Claudio Gay sequence (Mpodozis and Clavero, 2002), which is well exposed in the Cordillera Claudia Gay to the west (Stop 6-8). These sediments and volcanic rocks were emplaced in the backarc of the late Oligocene to Miocene volcanic arc, which is preserved in the Maricunga Belt some 35 km to the west (Day 7). The Segerstrom lavas are generally olivine- and plagioclase-bearing mafic flows (48%–57% SiO2) that dip gently eastward and have K-Ar ages of 25–24 Ma (Mpodozis et al., 1996; Kay et al., 1999). Their tholeiitic (FeO/MgO = 1.0–2.4) affinities, moderate (2%–3%) K2O contents, and low La/Ta (24–26) and Ba/La (16–19) ratios are consistent with eruption in a mildly extensional backarc. Their REE patterns (La/Yb = 9–14) are consistent with the parent lavas being small degree melts of mantle peridotite. Decreasing εNd (+4.3 to +1.5) and increasing 87Sr/ 86Sr (0.7039–0.7033) values show that the parent magmas were contaminated in the crust. The Segerstrom belt is one of several occurrences of Oligocene and early Miocene backarc mafic lavas in the central Andes. Others are the Máquinas basalt in the Chilean flat slab (30°S; Kay et al., 1991) and the Chiar Kkollu basalts in the Bolivian Altiplano (Davidson and de Silva, 1992). Isotopic differences between the Segerstrom basalts and the Central Volcanic Zone Incahuasi and San Francisco mafic flows (87Sr/86Sr ~0.7055; εNd ~+2) at Stop 6-1 reflect a change from a less enriched early Miocene backarc mantle to a more enriched Quaternary mantle wedge. This change is consistent with incorporation of crustal components into the mantle wedge by intense forearc subduction erosion at the time of arc migration (Kay and Mpodozis, 2000; Kay, 2006). The Cerros Amarillos lavas seen to the northeast are one of the oldest manifestations of arc-like andesitic to dacitic volcanism in the backarc of the Maricunga Belt arc. A dacite sample from a dome in this group yielded a K/Ar age of 13.9 ± 0.7 Ma (Mpodozis et al., 1996). This age fits with the middle Miocene expansion of arc volcanism into the backarc of the southern Puna as indicated by the abundant Miocene volcanic rocks in that region. The two-pyroxene andesite lavas (61% SiO2) from the Pliocene Rodrigo volcano to the south are distinctive for their glassy texture and chemical similarities to the southern Puna glassy lavas. The Rodrigo lavas have a K/Ar age of 4.4 ± 0.6 Ma (Mpodozis et al., 1996; Fig. 6.5). Chemical similarities of these lavas with the southern Puna lavas include pronounced arc-like HFSE depletions (La/Ta = 70–80) and steep REE patterns (La/Yb ~50). Volcanic rocks like these are largely restricted to the latest Miocene and Pliocene and reflect the special magmatic conditions at the time of arc migration (see Fig. 6.1; Kay and Mpodozis, 2002).

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Stop 6-7: Nevado de Tres Cruces Complex and Campo de Piedra Pomez Directions: Near 27°03′38″S; 68°55′00″W; 4350 m asl. The Nevado de Tres Cruces Massif seen to the east and south at this stop is formed by three large dacitic cones—Norte in the north (6206 m), Centro in the center (6629 m), and Sur in the south (6748 m), aligned along a NNW-trending regional fault system. The southern cone is considered to be the source of the nearly 100 m thick dacitic Tres Cruces ignimbrite (Pampa Blanca in Fig. 6.7, conspicuous white ignimbrite on TM scene in Fig. 6.4) with a 40Ar/39Ar age of 0.067 Ma (Gardeweg et al., 2000). The lavas, domes, and ignimbrites of the Macizo Tres Cruces volcanic complex (Fig. 6.7) are younger than 1.5 Ma. The youngest 40 Ar/39Ar age (0.28 ± 0.022 Ma) is from the southern cone. The dacitic pumice of the Campo de Piedra Pomez near the road at this stop has a 40Ar/39Ar age of 1.56 ± 0.15 Ma. Older Pliocene dacitic

flows around the Nevado de Tres Cruces Massif include the La Barra flow (2.3 ± 0.3 Ma) to the west (Fig. 6.5), the Lemp (2.8 ± 0.3 Ma), and Cristi (2.5 ± 1.3 Ma) volcanoes to the north, and the Cerro El Plateado flow to the east (Gardeweg et al., 1997, 2000). Stop 6-8: Rio Lamas—Middle Miocene Syntectonic Strata and Interbedded Ignimbrite Sheets in Cordillera Claudio Gay Directions: Watch for turnoff to Rio Lamas Waterfall from Route 31 near 27°04′54″S; 68°55′ (there are several tracks). Stop is along road, northeast of waterfall at 27°04′54″S; 68°55′59″W; 4280 m asl. After this stop, take Route 31 west. The route heads north along the Salar de Maricunga in the pre-Andean depression. At the north end of the salar, take the main road north to Chilean Customs and Immigration, and then continue north to La Ola. From La Ola, the road passes the south side of the Salar de

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Map of TRES CRUCES Region by Gardeweg et al. (2000) morraines glaciers volcanic vents normal faults, down side indicated Argentine-Chilean border K/Ar age Ar/Ar age

Figure 6.7. Geologic map of the Tres Cruces region after Gardeweg et al. (2000). Ages are from Gardeweg et al. (2000) and González-Ferrán et al. (1985). The location of Stop 6-7 on the Llano de Tres Cruces where the Tres Cruces ignimbrite is exposed is indicated.

Neogene evolution: central Andean Puna Plateau and southern-central volcanic zone Pedernales and heads to Montandón. West of Montandón, take the paved road north to Portal del Inca and continue to the town of El Salvador, where there are restaurants and hotels. The distance from the turnoff from Route 31 to the paved road to Porta del Inca is ~100 km. The distance along paved road to El Salvador is ~32 km. This stop at the southern termination of the Cordillera Claudio Gay provides a view of the synorogenic gravels and interbedded ignimbrites of the middle Miocene Rio Lamas sequence. Regionally, the Cordillera Claudio Gay (Figs. 6.1, 6.4, and 6.5) is an almost 150-km–long, north-south range with peaks up to 4800 m. At the latitude of Stop 6-8, the Cordillera Claudio Gay forms the eastern boundary of the modern pre-Andean depression that hosts the active, internally drained Salar de Maricunga and Salar de Pedernales basins. The western boundary of the preAndean depression is the Maricunga Belt, some 20 km to the west. The Maricunga Belt and the Salar de Maricunga will be well seen as the route of the trip heads west descending the western slope of the Cordillera de Claudia Gay. The Cordillera de Claudia Gay exposes a more than 1.5-km– thick Eocene to middle Miocene volcanic and sedimentary sequence that overlies Late Paleozoic acidic volcanic and granitoid basement (Mpodozis and Clavero, 2002). The oldest Tertiary unit is the Claudio Gay sequence, which is mainly composed of coarse dacitic pyroclastic deposits and domes. Epiclastic sediments and saline-lake carbonates accumulated in small intramontane basins from the coeval Rio Juncalito sequence (Mpodozis and Clavero, 2002). The Claudio Gay sequence is capped by the late Oligocene–early Miocene backarc Segerstrom basalt (Stop 6-6). At ca. 20–21 Ma, this depositional regime changed dramatically as east-verging, high-angle reverse faults uplifted the Late Paleozoic basement of the Cordillera Claudia Gay. Intense volcanism followed as large 20–19 Ma dome complexes with extensive block and ash-flow aprons erupted along the northern Cordillera Claudio Gay. These domes are covered by the widespread 18–19 Ma dacitic Vega Helada ignimbrite that has been correlated with the huge Rio Frio ignimbrite (Cornejo and Mpodozis, 1996). These ignimbrites are likely linked with the collapse that created the giant Aguilar caldera ~79 km to the north (e.g., Mpodozis and Clavero, 2002). In the middle Miocene, the Cordillera de Claudia Gay was affected by the last compressional deformation in the region of the modern pre-Andean depression. Evidence for this deformation comes from the alluvial gravels interbedded with distal ignimbrites (K/Ar age of 15–16 Ma) in the Rio Lamas sequence. These gravels show progressive unconformities and intraformational folds indicative of synsedimentary deformation (Gardeweg et al., 1997; Mpodozis and Clavero, 2002). DAY 7—THE MARICUNGA BELT: THE MIOCENE ARC FRONT The objective of Day 7 is to examine the Tertiary volcanic units of the Maricunga Belt where the active arc front was located from the late Oligocene to the late Miocene (26–6 Ma, Fig. 7.1). Stop 7-1 in the El Salvador mining district provides a view of

69°15′ W

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Dona Ines 20 km Salar de Pedernales

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7-2

Paleozoic granite/ rhyolite 26°30′ S

ESPERANZA

Barros Negros

La Ola

7-3 Cerro Bravos los Corrales La Coipa

Juncalito La Sal

Salar de Maricunga

Maricunga

7-4

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27°00′ S

Figure 7.1. Thematic Mapper image of the northern part of the Maricunga Belt showing Stops 7-2 to 7-4. Stop 7-1 is off the image to the west, and Stop 7-2 is near the edge.

pre-Maricunga Belt Cenozoic magmatism in the Chilean Precordillera just east of the Maricunga Belt (Figs. 7.1 and 7.2). Stop 7-2 examines the Miocene Atacama gravels associated with the uplift of the Puna plateau. Stops 7-3 to 7-6 provide representative views of the early Miocene, middle Miocene, and late Miocene magmatic styles of the Maricunga Belt. Stop 7-1: Paleocene El Salvador Caldera and El Salvador Porphyry Copper Deposit Directions: Take paved road south from El Salvador and stop along road ~3.5 km south near 26°17′34″S; 69°34′27″W; 2490 m asl. The area around the El Salvador mining district provides an excellent opportunity to examine the earliest Cenozoic stages of pre-Maricunga Belt magmatism and associated mineralization along the western Andean slope in the Chilean Precordillera.

Neogene evolution: central Andean Puna Plateau and southern-central volcanic zone Figure 7.2. Map of the northern Maricunga Belt and surrounding region from Mpodozis et al. (1995) showing locations of Stops 7-1 to 7-4 relative to the geology of the region. Compare with Thematic Mapper image in Figure 7.1. Map symbols are Pz (Paleozoic units), Mz (Mesozoic units), K-Ti (Upper Cretaceous to Paleocene units, ca. 70–50 Ma); Cerro Bravos–Barros Negros Group: eMi (early Miocene initial ignimbrites), eMa (early Miocene andesites), eMd (early Miocene domes), eMt (early Miocene); Maricunga Group: mMl (middle Miocene lavas), mMd (middle Miocene domes), mMt (middle Miocene tuffs).

Cenozoic volcanism in this region, as elsewhere in the Central depression and Precordillera of northern Chile began in the Paleocene when a series of stratovolcanoes, collapse calderas, and rhyolitic dome complexes erupted over deformed Jurassic to Late Cretaceous volcanic and sedimentary rocks (Mpodozis and Ramos, 1990; Cornejo et al., 1993a, 1993b, 1997, 1999; Arriagada et al., 2006). One of the largest of these centers was the 61–57 Ma El Salvador caldera seen from this stop. This center erupted high-K, calc-alkaline trachybasalts and trachyandesites and sanidine-bearing rhyolitic lavas, domes, and tuffs in response to the collapse of a shallow-level, zoned magma chamber. The lavas and ignimbrites have within-plate trace element signatures consistent with eruption in a period of very slow or arrested Farallon plate subduction. REE patterns show the magmas equilibrated with residual amphibole in a moderately thick crust (Cornejo et al., 1993b, 1999; Cornejo and Matthews, 2001). During a time of highly oblique, fast convergence between the Farallon and South American plates in the Eocene–early Oligocene (45–35 Ma) (Pardo Casas and Molnar, 1987), the Incaic deformation affected large parts of the Chilean Precordillera and formed the more than 1000-km–long Domeyko fault system (Tomlinson et al., 1994; Cornejo et al., 1997; Maksaev and Zentilli, 1999). The major fault in the El Salvador region is the left-lateral strike-slip Sierra del Castillo fault (Fig. 7.2), which will be crossed on the route to Stop 7-2. This fault separates Mesozoic volcanic rocks on the west from Jurassic–early Cretaceous backarc basin marine sediments in the Tarapacá backarc basin on the east side (Mpodozis and Ramos, 1990; Cornejo et al., 1993a, 1998). The spectacular thin-skinned folds and thrusts developed in the dominantly Mesozoic carbonate sequences of the Potrerillos fold-thrust belt are beautifully exposed along the deep canyons of the Quebrada Asientos and Río de la Sal. This thrust belt is seen on the route to Stop 7-2 and as the road continues eastward through Montandón. Ages on syn- and post-tectonic intrusions along the Domeyko fault system show that the Incaic deformation had begun in the El Salvador region by 42 Ma, was occurring at 36 Ma, and was over by 32 Ma (Tomlinson et al., 1994; Cornejo et al., 1993a). This deformation produced important thickening and exhumation and uplift of the crust as shown by apatite fission-track ages of 40–38 Ma (Maksaev and Zentilli, 1999; Nalpas et al., 2005). Extreme transpressive conditions trapped most of the magma in the crust. Small subvolcanic stocks and

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porphyries emplaced along major fault strands are hosts to giant Eocene-Oligocene porphyry copper deposits like those at Chuquicamata and La Escondida to the north. The El Salvador porphyry copper, which is often considered the type example of Chilean porphyry copper deposit (Cornejo et al., 1997), is hosted in the large Indio Muerto rhyolitic dome (58 ± 2 Ma, sensitive high-resolution ion microprobe [SHRIMP] U-Pb age; Cornejo et al., 1997, 1999) along the rim of the Paleocene El Salvador caldera. The chemistry of the El Salvador porphyry fits with mineralization being associated with fluids expelled from water-rich magmas as tectonic thickening led to destabilization of amphibole-bearing lower crust and formed an eclogitic crustal keel in the Oligocene (Cornejo et al., 1999; Cornejo and Matthews, 2001). Stop 7-2: Atacama Gravels, San Andreas Ignimbrite, and Uplift of the Andean Plateau Directions: Continue south and east along the paved and major gravel roads toward the village of Montandón. Stop ~5.6–6 km east of the junction of the paved and gravel road on the road to Montandón near 26°21′45″S; 69°27′34″W; 2490 m asl. The voluminous Miocene Atacama gravels (Mortimer, 1973), which occur along the western slope of the Andes, can be seen from this stop. These unconsolidated gravels, whose thicknesses exceed 400 m in places, filled drainage networks and covered the pre-middle Miocene topography. Those at this stop filled the paleovalley of the Quebrada Asientos; similar sequences filled the paleovalley of the Rió de la Sal to the north (Fig. 7.2). The gravels were shed westward from the rising Puna Plateau to the east from ca. 18–11 Ma (Cornejo et al., 1993a; Nalpas et al., 2005) as the hyperarid climate developed that has prevailed in the Atacama Desert since the Miocene (Alpers and Brimhall, 1988; Rech et al., 2006; Nalpas et al., 2005). 40Ar/39Ar ages on interbedded ashes in the gravels range from 15.3 to 10.5 Ma (Owen et al., 2003). A regional pediment surface (Atacama Pediplain of Mortimer, 1973) developed as gravel accumulation ceased. Dating of cobbles on this surface using cosmogenic nuclides (10Be, 26Al, and 21Ne) yield exposure ages of 9 Ma (Nishiizumi et al., 2005). Near El Salvador, the pediment is covered by the widespread, upper Miocene San Andrés ignimbrite with biotite K/Ar ages ranging from 10.2 to 9.3 Ma (Clark et al., 1967; Mortimer, 1973; Cornejo et al., 1993a). The San Andrés ignimbrite is contemporaneous with Copiapó complex ignimbrites in the Maricunga Belt that will be seen at Stop 7-6. A remnant of the San Andrés ignimbrite on the pediment surface northeast of the abandoned Potrerillos airport can be seen from this stop. Introduction to Stops 7-3 to 7-6: Oligocene to Miocene Maricunga Belt Volcanic Arc Front The objective of Stops 7-3 to 7-6 is to examine the late Oligocene to late Miocene magmatic and tectonic history of the Maricunga Belt arc. The volcanism in the Maricunga Belt

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coincides with the shallowing of the subducting Nazca slab under the Chilean flat slab to the south (see Kay et al., 1991) and the uplift of the Puna plateau to the east. The magmatic history reflects crustal thickening as the subducting plate changed shape and modified the overlying crust and lithospheric mantle. Overviews of the Maricunga Belt can be found in Mpodozis et al. (1995) and Kay et al. (1994b, 1999). Other discussions of the Maricunga Belt region and individual volcanic centers are in Mercado (1982), González-Ferrán et al. (1985), Baker et al. (1987), Davidson and Mpodozis (1991), Walker et al. (1991), Mpodozis et al. (1991), Vila and Sillitoe (1991), Vila et al. (1991), Moscoso et al. (1993), Cornejo et al. (1993a, 1998), McKee et al. (1994), Sillitoe et al. (1991), Tomlinson et al. (1999), and Kay and Mpodozis (2000, 2001, 2002). The petrology and geochemistry of the Maricunga Belt arc rocks are discussed by Kay et al. (1994b, 1999), Mpodozis et al. (1995), and Kay and Mpodozis (2000, 2002). The volcanic rocks are mostly medium- to high-K calc-alkaline andesites and dacites (56%–68% SiO2) whose geochemistry reflects a temporal trend from low-pressure, pyroxene-dominated to higher pressure, garnet-bearing eclogitic residual mineral assemblages in equilibrium with the magmas. Increasingly steeper heavy REE patterns show a growing role for garnet, and decreasing negative Eu anomalies and increasing Sr and Na2O concentrations show a decreasing role for plagioclase. The trends are interpreted to reflect the temporal evolution of these magmas in relation to a cooling mantle wedge and a thickening crust. Extreme chemical characteristics in late Miocene to Pliocene lavas reflect magma evolution in a very thick crust (>65 km) and introduction of continental crust into the mantle wedge by forearc subduction erosion as the arc front was displaced eastward. The magmatic and tectonic history of the Maricunga Belt can be summarized in five stages:

dacitic stratovolcanic complexes along the length of the Maricunga Belt. From north to south, these centers include the Doña Inés, Ojos de Maricunga (Stop 7-4), Santa Rosa, Pastillos-Pastillitos (Stop 7-5), La Laguna, and CadillalYeguas Heladas volcanoes. Ignimbrites were emplaced during the initial stages at some centers. Most centers are little eroded and preserve much of their original form. The Atacama gravels seen at Stop 7-2 were deposited in the Precordillera and Central depression to the west at this time. The third stage ended with the emplacement of structurally controlled, shallow-level, quartz-dioritic stocks hosting gold and copper mineralization (Marte, Lobo, and AldebaránCasale gold porphyries; Vila and Sillitoe, 1991). 4. Late Miocene (11–10–7 Ma): A radical change in the distribution of volcanic centers occurred during the fourth stage as most of the volcanic activity north of 27.5°S became concentrated in the silicic andesitic to dacitic Copiapó volcanic complex (Stop 7-6). 5. Late Miocene to Pliocene (7 Ma to 5 Ma): The last stage is marked by the eruption of small-volume bimodal centers at the southern end of the Maricunga Belt. These eruptions include ignimbrites associated with the fault-controlled Jotabeche rhyodacitic caldera and glassy mafic andesitic to andesitic Pircas Negras lavas emplaced along faults (Mpodozis et al., 1991, 1995; Kay et al., 1991, 1994b). The chemistry (very high La/Yb ratios and Na2O and Sr contents) of these magmas are uncommon in the central Andes and have been associated with a very thick crust and contamination of the mantle wedge by crustal material removed from the forearc by subduction erosion (e.g., Kay and Mpodozis, 2002).

1. Late Oligocene to early Miocene (26 to 21–20 Ma). During the first stage, andesitic to dacitic volcanic complexes erupted over a moderately dipping subduction zone through about a 40-km–thick crust. Centers include the large andesitic to dacitic Cerro Bravos–Barros Negros stratovolcano and dome complexes in the north (Stop 7-3) and widespread, small multiple dacitic dome complexes in the south. Many of the dome complexes and associated tuff and pyroclastic breccia rings are hydrothermally altered. Epithermal, high-sulfidation, gold-silver and porphyry-style gold mineralization occurred during this stage at La Coipa, Esperanza, Pantanillo, Refugio, and La Pepa (Vila and Sillitoe, 1991; Mpodozis et al., 1995; Muntean and Einaudi, 2001). 2. Early Miocene (21–17 Ma). The second stage was characterized by a virtual volcanic lull during a period of compressional deformation and crustal thickening. Evidence for compressional deformation was seen in the Cordillera Claudio Gay at Stop 6-8 (Mpodozis and Clavero, 2002). 3. Middle Miocene (17–12–11 Ma). The third stage was marked by the construction of voluminous, andesitic to

Directions: From Stop 7-2 continue on major gravel road through Montandón to La Ola. At La Ola, take major road to south and stop near 26°40′40″S; 69°03′44″W; 3758 m asl. The overlapping Cerros Bravos–Barros Negros volcanoes seen to the west from this stop (Figs. 7.1 and 7.2) were emplaced in the earliest volcanic episode of the Maricunga Belt. They erupted over a moderately dipping subduction zone through a crust that was likely near 40 km thick. These centers mark the frontal arc west of the backarc Segerstrom basalt seen at Stop 6-6. The Cerros Bravos–Barros Negros centers are located along sinistral northwest-trending strike-slip faults that are part of the Eocene Potrerillos–Maricunga system that parallels the southern side of the Paleozoic Pedernales batholith (Cornejo et al., 1993a; Mpodozis et al., 1995; Tomlinson et al., 1999). K/Ar ages from these centers range from 25 ± 1 Ma to 21.7 ± 1 Ma (Moscoso et al., 1993; Cornejo et al. 1993a; McKee et al., 1994; Kay et al., 1994b; Mpodozis et al., 1995; Tomlinson et al., 1999). Activity at the Cerros Bravos–Barros Negros centers began with the eruption of small-volume rhyodacitic ignimbrites (eMi in Fig. 7.2) that were subsequently covered by main stage

Stop 7-3: Late Oligocene–Early Miocene Cerros Bravos– Barros Negros Volcanic Complexes

Neogene evolution: central Andean Puna Plateau and southern-central volcanic zone pyroxene- and hornblende-andesitic lava and block and ash deposits (eMa in Fig. 7.2). Crystal-rich hornblende and biotitebearing dacitic domes (eMd in Fig. 7.2) were then emplaced in the cores of the stratovolcanoes. A series of subcircular domes that host the Esperanza epithermal gold and silver deposit (Vila and Sillitoe, 1991; Moscoso et al., 1993) were emplaced in an 8-km–long belt along a reactivated WNW-trending fault zone (Cornejo et al. 1993a) on the northeastern flank of Cerros Bravos. The domes at Esperanza are surrounded by plinian rhyolitic lapilli tuffs (eMt in Fig. 7.2) with K/Ar ages ranging from 24 to 20 Ma. Alteration ages range from 23.2 ± 1.4 to 19.3 ± 0.7 Ma (Sillitoe et al. 1991; Moscoso et al. 1993). The northeasternmost domes are intruded by unaltered dacitic porphyry dikes with biotite K/Ar ages of 22.5 ± 1.2 and 22.4 ± 0.7 Ma (Cornejo et al., 1993a; Kay et al., 1994b). A few kilometers south of Cerro Bravos, a multistage dome complex was emplaced at La Coipa. The La Coipa complex is located where west-verging, north-trending thrusts (Vegas la Junta) intersect the sinistral northwest-trending Quebrada Indagua fault. The La Coipa dome cluster consists of five 300- to 400-m–wide, crystal-rich, dacitic to rhyodacitic domes. K/Ar ages from unaltered early Miocene domes range from 24.6 ± 0.9 Ma to 22.9 ± 0.8 Ma (Zentilli 1974; Moscoso et al., 1993); ages of hydrothermally altered domes range from 20.2 ± 1.2 to 17.3 ± 1.0 Ma (Sillitoe et al., 1991; Moscoso et al., 1993; Mpodozis et al., 1995). The domes are surrounded by an extensive coeval blanket of intensely altered, coarse pyroclastic breccias, and poorly welded lapilli tuffs with biotite K/Ar ages of 24.7 ± 0.7 and 24.0 ± 1.0 Ma. These rocks host the epithermal silver-gold mineralization at the La Coipa, Can-Can, and Purén mines and prospects (e.g., Cornejo et al., 1993a; Mpodozis et al., 1995). Volcanism in the La Coipa region ended with the emplacement of middle Miocene domes (mMd in Fig. 7.2; Mpodozis et al., 1995). Stop 7-4: Early Middle Miocene Volcanism— Ojos de Maricunga Stratovolcano Directions: Continue south, passing the Chilean Customs, and take a major gravel road to the southeast toward Laguna Santa Rosa; stop near 27°00′36″S; 69°03′56″W; 3785 m asl. The largest group of middle Miocene volcanic centers in the Maricunga Belt is the cluster of stratovolcanoes to the west and south of the Salar de Maricunga. The northernmost of these centers is seen to the west from this stop (Figs. 7.1 and 7.2). This is the well-preserved Ojos de Maricunga volcano (4985 m) with a basal diameter of 15 km and a central crater filled by a dacite dome dated at 15.8 ± 0.9 Ma (whole-rock K/Ar; Mpodozis et al., 1995). The slopes of the volcano are covered by unconsolidated hornblende andesite block and ash deposits (mMl on Fig. 7.2) that have yielded K/Ar ages from 16.2 ± 0.6 to 15.1 ± 0.7 Ma (Zentilli, 1974; Mpodozis et al., 1995) and a 40Ar/39Ar age of 14.5 ± 0.1 Ma (McKee et al., 1994). The block and ash deposits overlie two ignimbrites of uncertain origin (mMt on Fig. 7.2). The older is a welded red tuff

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(60%–62% SiO2) that is up to 100 m thick and has a wholerock K/Ar age of 15.8 ± 0.8 Ma. The younger, which is exposed on the southwestern slope, is a slightly welded, pumice-rich biotite-bearing ignimbrite with biotite K/Ar ages of 14.3 ± 1.6 Ma and 13.7 ± 2.6 Ma (see Zentilli, 1974). Other middle Miocene volcanic centers to the south are principally made of hornblende- and pyroxene-bearing andesite. They include the Santa Rosa, Cerro Lagunillas, Pastillitos, and Cerro Las Cluecas volcanoes (Figs. 7.4 and 7.5). Blocks from a coarse blanket of reworked pyroclastic block and ash deposits on the slopes of the Santa Rosa cone have K/Ar ages of 15.4 ± 0.55 Ma (hornblende, McKee et al., 1994) and 13.8 ± 0.6 Ma (whole rock, González-Ferrán et al., 1985). Whole-rock K/Ar ages from a Cerro Lagunillas lava, a block from the Pastillitos center, and a Cerro Las Cluecas lava range from 16.2 ± 0.6 to 15.9 ± 1.4 Ma (Mpodozis et al., 1995). Stop 7-5: Earliest Late Miocene Volcanism and the Gold Porphyries: Pastillos Volcano and Mina Marte Directions: Continue south, passing the road to Laguna Santa Rosa and Tambo and the Ex Mina Marte. South of the Ex Mina Marte, stop on the road near 27°12′04″S; 69°00′43″W; 4000 m asl. Just to the east of the Ojos de Maricunga centers are the Villalobos and Pastillos volcanoes. Their K/Ar ages are summarized by Mpodozis et al. (1995) (Fig. 7.3). The Villalobos center is a deeply eroded stratovolcano with ages of 16.1 ± 2.1 and 14.8 ± 0.8 Ma (K/Ar) and 13.3 ± 0.5 Ma (40Ar/39Ar, McKee et al., 1994). The andesitic to dacitic Pastillos volcano is superimposed on the Villalobos volcano. Flows on the western flank of the Pastillos volcano have ages from 12.0 ± 1.8 to 12.9 ± 0.51 Ma. The Pastillos lavas and flows west of Cerro Las Cluecas (K/Ar ages of 13.1 ± 1.6–12.4 ± 1.4 Ma) are the youngest in the Ojos de Maricunga region. The Pastillos-Villalobos center is intruded by shallow level, hornblende-bearing, andesitic and dacitic to quartz diorite porphyries. Those hosting gold-rich quartz stockwork alternation in the Marte, Lobo, Escondido, and Valy prospects along the eastern flank of the Villalobos center are called gold porphyries (Sillitoe, 1994; Mpodozis et al., 1995). The main targets are in the underlying Oligocene tuffs and the core of Pastillos volcano. The gold porphyries were emplaced along a west-northwest– trending fault system that projects from the Valle Ancho to the east into the northwest-trending sinistral Eocene fault system of the Porterillos–La Coipa region to the west (Tomlinson et al., 1994). An unusual association of porphyry-style mineralization and epithermal advanced argillic alteration (telescoping) in the Marte-Lobo district was attributed to rapid erosion linked to volcanic edifice collapse as the porphyry intruded by Sillitoe (1994). A problem is that physical evidence for collapse at Marte is unclear (Mpodozis et al., 1995). The time of mineralization is based on alunite K/Ar ages of 13.3 ± 0.42 Ma (Sillitoe et al., 1991) and 12.0 ± 0.69 Ma (Zentilli et al., 1991)

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Figure 7.3. Graph showing peaks of K/Ar and Ar/Ar ages in the Maricunga Belt. Figure is from Kay et al. (1994b).

in acid-sulfate altered rocks and a K/Ar whole-rock age of 11.4 ± 0.5 Ma on the Valy porphyry (see Mpodozis et al., 1995). The first Maricunga Belt deposit called a porphyry gold deposit was at the Marte prospect (Vila et al., 1991). This hornblende-biotite diorite porphyry stock was emplaced and mineralized at ca. 13–14 Ma. The stock is in a 2-km–diameter subcircular depression that is considered to be the small eroded summit crater of the Pastillos stratovolcano (Sillitoe et al., 1991; Mpodozis et al., 1995). It consists of two mineralized intrusives and a younger weakly porphyritic microdiorite. The mineralization is in a stockwork of quartz veinlets surrounded by chlorite-sericiteclay alteration, accompanied by disseminated pyrite and hematite that likely formed by hypogene martitization (replacement of magnetite by hematite). Isolated remnants of hydrothermal biotite and alkali feldspar indicate an overprint of potassic alteration. The deposit contains ~66 tons of gold. An open-pit, heapleaching operation began in 1989, but closed shortly after, due to metallurgical problems. Stop 7-6: Late Miocene Copiapó Volcanic Complex About 0.6 km south of Stop 7-5, turn west on a track south of a lagoon; continue to near 27°13′17″S; 69°04′56″W; 4130 m asl. This road is best traversed with four-wheel drive. Late Miocene volcanic activity in the Maricunga Belt from 11 to 7 Ma was largely restricted to the Copiapó volcanic complex at the intersection of the northwest-trending Valle Ancho– Potrerillos fault system with the Maricunga Belt. The silicic andesitic to dacitic pyroclastic flows, domes, and lavas that make up the Copiapó complex cover an area of more than 200 km2 The Copiapó complex is shown on the satellite image in Figure 7.4 and the map in Figure 7.5 from Mpodozis et al. (1995). Volcanic activity took place in the two main stages discussed below. Ages are summarized in Mpodozis et al. (1995) and Kay et al. (1994b).

Stage 1 (11–10 Ma) Volcanic activity in the Copiapó complex began with the construction of the Azufre stratovolcano, north of the Laguna Negro Francisco. The main edifice is made of hornblende dacite flows and block and ash-flow deposits that are intruded by late-stage, fine-grained dacitic porphyries. Extensive hydrothermal alteration accompanied emplacement of the porphyries. The northern side of the Azufre center is covered by flows from the Azufre Norte center. Whole-rock K/Ar ages are 11.2 ± 0.9 and 9.2 ± 0.9 Ma for the Azufre cone, 11.9 ± 0.7 and 10.5 ± 0.5 Ma for the Azufre Norte flows, and 10.8 ± 0.5 Ma for the Azufre porphyries. At about the same time, large explosive eruptions likely linked to caldera formation produced the well-exposed pyroclastic deposits north and west of the main Copiapó center. These 100- to 300-m–thick deposits are visible from this stop. A basal sequence consists of smallvolume ignimbritic flows and block and ash deposits alternating with thin plinian ash falls and proximal detrital deposits, lahars, and fluvial sandstones. They are overlain by the lightgray Copiapó I ignimbrite, which is a small-volume, weakly to moderately welded unit with abundant pumice and lithic fragments. The younger and overlying pink to reddish-brown Copiapó II ignimbrite is a welded columnar-jointed cooling unit with concentrations of large pumice fragments at the top. Both ignimbrites are hornblende-bearing silicic andesite and/or dacite. Ignimbrite whole-rock K/Ar ages range from 11.0 ± 2.2 Ma to 9.6 ± 0.8 Ma. To the south and east, the Copiapó ignimbrites interfinger with a large coalesced dome and lava-dome complex that is in turn intruded by two fine-grained porphyritic stocks with acidsulfate hydrothermal alteration (Azufrera de Copiapó). Four whole-rock K/Ar ages from the dome complex range from 11.7 ± 0.6 to 9.8 ± 0.7 Ma. A porphyritic stock in the Azufrera has a K/Ar age of 10.8 ± 0.5 Ma.

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

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Figure 7.4. Thematic Mapper satellite image showing the locations of Stops 7-5 and 7-6 relative to the early Miocene Ojos de Maricunga, the middle Miocene Santa Rosa volcano, the late middle to earliest late Miocene volcanic centers hosting the gold porphyries (Marte, Lobo, Soledad), and the late Miocene Copiapó Volcanic Complex. The northwest-trending Valle Ancho is seen in the southeastern part of the image. This valley coincides with the Valle Ancho fault zone, which projects beneath the Copiapó center near the gold porphyries.

Stage 2 (8–7 Ma) The shield-like Copiapó dome complex produced in Stage 1 forms the base for the younger stage 2 Copiapó cone. The central edifice of the younger cone is composed of hornblende and biotite-bearing dacite lavas and lava domes. Four wholerock K/Ar ages range from 8.4 ± 0.4 Ma to 6.9 ± 1.1 Ma. A small parasitic dome at Cerro San Román, north of Copiapó has

a K/Ar age of 6.9 ± 1.6 Ma and a 40Ar/39Ar hornblende age of 6.6 ± 0.3 Ma (Mulja, 1986). Directions: End of field trip. To reach the city of Copiapó, return to the gravel road heading north. Continue through Ex Mina Marte and turn west on the road to Tambo and Laguna Santa Rosa. Join Route 31 at La Puerta and continue through La Puerta and Puquois to intersect the paved road to the city of

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Paleozoic to 69°10′ W Miocene basement

ash, block & ash

7-5 ig1 ig1 ig1

27°15′ S

7-6

ig2

27°15′ S

altered ig2 flows

dome 1 dome 2 ash, block&ash

altered

dacitic porphry lavas

Paleozoic to Miocene basement

gravels Figure 7.5. Geologic map from Mpodozis et al. (1995) showing volcanic units of the Copiapó volcanic complex and surrounding area and the locations of Stops 7-5 and 7-6. Units are discussed at Stop 7-6. Compare with Thematic Mapper Image in Figure 7.4.

Copiapó. The distance from the turnoff to Laguna Santa Rosa to Copiapó is ~190 km. An alternative and better maintained route is to continue north from Ex Mina Marte to the intersection of Route 31 and follow this road west to Copiapó. ACKNOWLEDGMENTS We would like to thank the participants of the Geological Society of America–Asociación Geológica Argentina “Backbone of the

Americas” field trip 405 that was held 9–15 April 2006 for their enthusiastic participation and their comments that improved the presentation of this guide. We would particularly like to thank Pablo Caffe, Magdalena Koukharsky, and Alejandro Pérez for their assistance in leading the trip and their invaluable logistical support. Alan Glazner, Robert Kay, and Víctor Ramos provided valuable comments that improved the presentation of the guide. This field trip guide is dedicated to the memory of William Hiss, who died in Antofagasta de la Sierra during the field trip in 2006.

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