Contents Dedication: Pablo Groeber (1885–1964) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . v Preface . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . vii 1. Overview of the tectonic evolution of the southern Central Andes of Mendoza and Neuquén (35 °–39°S latitude) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1 V.A. Ramos and S.M. Kay 2. Upper Cretaceous to Holocene magmatism and evidence for transient Miocene shallowing of the Andean subduction zone under the northern Neuquén Basin . . . . . . . . . . . . . 19 S.M. Kay, W.M. Burns, P. Copeland, and O. Mancilla 3. Deep seismic images of the Southern Andes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61 X. Yuan, G. Asch, K. Bataille, G. Bock, M. Bohm, H. Echtler, R. Kind, O. Oncken, and I. Wölbern 4. Neogene tectonic evolution of the Neuquén Andes western flank (37–39°S) . . . . . . . . . . . . . . . . 73 D. Melnick, M. Rosenau, A. Folguera, and H. Echtler 5. Intraplate deformation in the Neuquén Embayment . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 97 A. Mosquera and V.A. Ramos 6. Structural evolution and magmatic characteristics of the Agrio fold-and-thrust belt. . . . . . . . . 125 G. Zamora Valcarce, T. Zapata, D. del Pino, and A. Ansa 7. Synrift geometry of the Neuquén Basin in northeastern Neuquén Province, Argentina . . . . . . 147 E. Cristallini, G. Bottesi, A. Gavarrino, L. Rodríguez, R. Tomezzoli, and R. Comeron 8. The case for extensional tectonics in the Oligocene-Miocene Southern Andes as recorded in the Cura Mallín basin (36 °–38°S) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 163 W.M. Burns, T.E. Jordan, P. Copeland, and S.A. Kelley 9. Early to middle Miocene backarc magmas of the Neuquén Basin: Geochemical consequences of slab shallowing and the westward drift of South America . . . . . . . . . . . . . . . . 185 S.M. Kay and P. Copeland 10. Evolution of the late Miocene Chachahuén volcanic complex at 37°S over a transient shallow subduction zone under the Neuquén Andes . . . . . . . . . . . . . . . . . . . . . 215 S.M. Kay, O. Mancilla, and P. Copeland 11. Miocene to Quaternary deformation of the Guañacos fold-and-thrust belt in the Neuquén Andes between 37 ° and 37 °30’S . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 247 A. Folguera, V.A. Ramos, E.F. González Díaz, and R. Hermanns iii
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Contents 12. Late Cenozoic extension and the evolution of the Neuquén Andes . . . . . . . . . . . . . . . . . . . . . . . 267 A. Folguera, T. Zapata, and V.A. Ramos 13. Upper Pliocene to Lower Pleistocene volcanic complexes and Upper Neogene deformation in the south-central Andes (36°30’–38 °S). . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 287 F. Miranda, A. Folguera, P.R. Leal, J.A. Naranjo, and A. Pesce 14. The Pliocene to Quaternary narrowing of the Southern Andean volcanic arc between 37° and 41°S latitude. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 299 L.E. Lara and A. Folguera 15. Geochemistry and isotopic characteristics of the Caviahue-Copahue volcanic complex, Province of Neuquén, Argentina . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 317 J.C. Varekamp, J. Maarten deMoor, M.D. Merrill, A.S. Colvin, A.R. Goss, P.Z. Vroon, and D.R. Hilton Index. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 343
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Geological Society of America Special Papers Dedication Pablo Groeber (1885−1964) Geological Society of America Geological Society of America Special Papers 2006;407;v doi: 10.1130/0-8137-2407-4.v
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Dedication Pablo Groeber (1885–1964)
The outstanding characteristics of the Andes of Neuquén and Mendoza were studied for more than 50 years by Dr. Pablo Groeber, who, through his pioneering work, established the first understanding of the geology of this part of the Andes. This German geologist, who received his doctorate degree in geology at the University of München, Germany, in 1897, arrived in Argentina when he was 26 years old. Prior to his arrival in Argentina, he conducted two large expeditions to the central Tien Shan, where he spent more than two years establishing the basis of the modern structural and stratigraphic knowledge of this remote and, at this time, unknown region of central Asia. Soon after being hired by the Geological Survey of Argentina, Groeber began one of the most creative investigations of the Andes of Argentina and Chile. He mapped vast areas of the cordillera, surveying his own topography and producing structural sections and sketches of the most remote parts of the region. He set up much of the present stratigraphic framework of Neuquén and Mendoza, and documented the basic structures and magmatic cycles related to the deformation and uplift of the Cordillera de Los Andes. He also recognized the main depositional sequences, the different uplift pulses, and the migrations in volcanic activity. In those years before isotopic ages were available, he identified and assigned ages to a series of basaltic and andesitic volcanic episodes associated with different Cenozoic diastrophic phases. Groeber wrote fundamental papers that led the way to the present state of knowledge in the region. His publications on the high Cordillera de Los Andes of Mendoza and Neuquén, on the structural evolution of the Neuquén Basin, and his geologic maps along the 70°W meridian form the basis of the modern geology of the region. The reader will find references to Groeber and mention of his seminal proposals in a number of chapters in this volume. During his tenure with the survey, Groeber also was teaching in the Universities of Buenos Aires and La Plata, where he strongly influenced several generations of geologists and supervised several doctoral theses. For a number of decades, his work was the main reference on the geological setting and evolution of the south central Andes. His disciples continued his work, improved his schemes, and refined the correlations of the geologic units, but the main tectonic framework that he proposed in the 1950s is still the core of our understanding of this sector of the Andes. v
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Geological Society of America Special Papers Preface Suzanne Mahlburg Kay and Víctor A. Ramos Geological Society of America Special Papers 2006;407;vii-x doi: 10.1130/0-8137-2407-4.vii
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Preface The Andes are the type example of an accretionary orogen that has developed without a continental collision. In an accretionary orogen, magmatism occurs over a subducting slab and is associated with episodes of contractional and extensional deformation that are explained by the relative rollback velocity of the subducting slab, episodes of shallowing and steepening of the subduction zone, and changes in convergence parameters. In this framework, the evolution of the northernmost Patagonian and southernmost Central Andean Cordillera between 35°S and 39°S along with the region to the east (retroarc) is a virtual endmember in understanding the development of an Andean-type accretionary orogeny. In contrast to the general conception of an Andean orogen, the Tertiary to Holocene evolution of this part of the Andes is generally characterized by basaltic to andesitic arc volcanic centers, relatively low relief, both periods of compression and extension, a relatively small amount of crustal shortening, and extensive retroarc mafic volcanism. This evolution stands in marked contrast to a typical Andean model that is strongly influenced by the Neogene history of the central most Andes; a region that constitutes only a fraction of the length of the long-lived orogen and a timeframe that constitutes only a part of its Jurassic to Holocene history. Unlike the southernmost Central Andes, the evolution of the central most Central Andes is marked by andesitic to dacitic stratovolcanoes and gigantic ignimbrites complexes, large amounts of contractional deformation east of the arc that have resulted in extensive crustal shortening, a relatively silicic crust whose mafic part could have been removed by delamination, and the dramatic uplift of one of the world’s largest and highest plateaus— the Puna-Altiplano. A major objective of this volume is to examine the Tertiary to Holocene tectonic and magmatic evolution from the arc to the retroarc in the distinctive and until recently little understood end-member of the Andean accretionary orogen between 35°S and 39°S. The Tertiary to Holocene evolution of the eastern slope and retroarc of the Andes in the Argentine provinces of Neuquén and Mendoza between 35°S and 39°S has been strongly influenced by the Mesozoic development of the Neuquén Basin, which is one of the most productive and historically important hydrocarbon basins in South America. The literature on the region has concentrated heavily on the late Triassic to Paleocene evolution of the basin as the subsequent history has traditionally been of less interest to the petroleum industry. The detailed studies of the region have emphasized the sedimentary aspects of the Triassic to Early Jurassic rifting history of the basin associated with the breakup of Pangea and its subsequent
Mesozoic evolution as an Andean foreland basin. Numerous other papers have dealt with the richness of the Mesozoic fossil assemblages that have yielded some of the world’s most spectacular ammonite, dinosaur and vertebrate fossils. This picture has changed in the past decade due to several initiatives financed by the petroleum industry and government funding agencies. The impetus has come from the role of postMesozoic deformation and magmatism in modifying the Mesozoic configuration of the Neuquén basin and in controlling petroleum migration. Another contributing factor has been that large numbers of oil wells have been drilled through postCretaceous magmatic flows. Questions have been posed concerning the style and effects of Tertiary to Holocene deformation, the age and origin of the extensive retroarc volcanic rocks, and the tectonic controls of deformation and magmatism. A paucity of post-Paleocene sedimentary strata has led to studies that emphasize magmatism and deformation as the keys to understanding the tectonic evolution of this region whose history reflects the development of the long-lived Andean margin. The objective of this volume is to present new data and ideas that have emerged relative to the post-Mesozoic evolution of the Neuquén Basin. These studies have led to major steps in understanding the arc and retroarc evolution of this distinctive segment of the Andes. The papers in this volume constitute a complement to the literature on the young volcanic centers of the Southern Volcanic Zone arc, and to a volume titled The Neuquén Basin: A case study in sequence stratigraphy and basin dynamics (published by the Geological Society of London). The papers in the later volume emphasize the Mesozoic stratigraphy, sedimentary geology, biostratigraphy, paleoecology, paleogeography and deformational history of the Neuquén Basin. Among the results in the chapters of this volume are major revisions in the timing of Mesozoic and Cenozoic magmatic and deformational events in the arc and retroarc, in models for relating the arc and retroarc magmatism and the extensional and contractional deformational history to plate convergence vectors, and in the importance of pre-Tertiary structures in controlling younger deformational styles. Proposals are advanced for transient Miocene shallowing of the Andean subduction zone in association with development of Laramide-style magmatism and uplift far east of the trench, and for an association of Pliocene to Quaternary arc to retroarc extension and retroarc mafic magmatism in the Payenia Large Igneous Province with steepening of the subducting slab over a hydrated mantle. The first chapter, by V.A. Ramos and S.M. Kay, presents an overview of the Mesozoic to Holocene tectonic evolution of vii
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the Andes between 35°S and 39°S with an emphasis on the southern Mendoza and Neuquén provinces of Argentina. The paper presents a synthesis of the current state of knowledge on the structural evolution of the region and draws attention to the importance of the control of older structural and lithospheric scale features, shifting convergence parameters, and changing slab configurations on the spatial and temporal evolution of the region. The second chapter, by S.M. Kay, W.M. Burns, P. Copeland, and O. Mancilla, is complementary to the first chapter in presenting a synthesis of the Upper Cretaceous to Holocene magmatic evolution of the Neuquén basin between 36°S and 38° S and in presenting the evidence for Kay’s original suggestion for a transient Miocene shallow subduction north of the Cortaderas lineament. A survey of the spatial and temporal distribution of the magmatic rocks along with new major and trace element analyses (90 samples), new isotopic data (12 samples), and 12 new 40Ar/ 39Ar ages are used to argue two major points. (1) The configuration of the subduction zone from the trench to the arc has been little modified since the Upper Cretaceous. (2) A transient period of shallow subduction that began at ca. 20 Ma and peaked in the late Miocene was followed by Pliocene steepening that led to eruption of widespread mafic magmas that form the backarc Payenia Large Igneous Province. The southern margin of the shallowly dipping subduction zone is argued to be near the Cortaderas lineament that essentially bounds the limit of Neogene retroarc magmatism. The third chapter, by X. Yuan, G. Asch, K. Bataille, and coworkers at the GeoForschungsZentrum in Potsdam, Germany, reports some of the initial results of the TIC-TAC project whose primary aim is a study from the Chile trench across the forearc to the main Andean cordillera at this latitude. The paper presents some of the first deep seismic images of the southern Andes between latitude 36° and 40°S. The results, which are from receiver function images, show the oceanic Moho of the subducted Nazca plate imaged down to a depth of ~100 km in good correspondence with results from Wadati-Benioff zone seismicity and wide-angle seismic reflections, the continental Moho at a depth of ~40 km beneath the Main Cordillera, and the eastward shallowing of the Moho to ~35 km beneath the western Neuquén basin. An intriguing result shows the continental Moho locally shallowing to an apparent depth of 30 km beneath the Loncopué graben on the eastern slope of the Andes where Pliocene to Holocene extension is concentrated. The fourth chapter, by D. Melnick, M. Rosenau, A. Folguera, and H. Echtler, uses new field observations and the published literature to discuss the Neogene evolution of the western flank of the Andes in Chile between 37°S to 39°S. This evolution is marked by: (1) Oligocene to middle Miocene extension followed by late Miocene shortening coincident with uplift, exhumation, inversion, and a volcanic gap in the Main Cordillera; (2) Pliocene to early Pleistocene extension with reestablishment of the arc and transtension along the intra-arc zone; and (3) late Pleistocene to Holocene narrowing of the arc and localized extension-transtension along the axial intra-arc
zone. The main uplift of the cordillera is placed between 11 and 6 Ma. The cessation of contraction is argued to be linked to an increase in slab angle that also triggered extension along the orogenic front and onset of arc-parallel strike-slip faulting. The crustal-scale dextral strike-slip Liquiñe-Ofqui fault zone, which concentrates deformation south of 38°S and helps to accommodate oblique subduction is reviewed and analyzed. The fifth chapter, by A. Mosquera and V.A. Ramos, presents a synthesis of intraplate deformation in the Neuquén Basin based on previously unpublished two- and three-dimensional seismic images. The discussion emphasizes: (1) the role of Paleozoic basement fabrics in the development of Mesozoic and Cenozoic deformational fabrics, (2) the important role of fabrics inherited from the Late Permian collision of an allochthonous Patagonian terrane whose suture with Gondwana is argued to be under the Neuquén Basin, (3) how the sequence and location of uplifts, inversion of half-graben systems, and strike-slip faults can be used to demonstrate a shifting spatial and temporal pattern of the main stresses across the region, and (4) how the changing deformational pattern can be linked to changes in the convergence parameters between the South America (Gondwana) and oceanic plates to the west. The sixth chapter, by G. Zamora Valcarce, T. Zapata, D. del Pino, and A. Ansa, shows that the Agrio fold-and-thrust belt in the retroarc between 37°S and 38°S was subjected to major contractional deformation in the Late Cretaceous and the Miocene. The paper presents a description of the Agrio belt, crucial new 40Ar/ 39Ar ages that for the first time show the importance of Late Cretaceous contractional deformation in the Agrio belt, and new geochemical data on the Cretaceous to Eocene volcanic rocks from which the 40Ar/ 39Ar ages were obtained. The seventh chapter, by E. Cristallini, G. Bottesi, A Gavarrino, L. Rodríquez, R. Tomezzoli, and R. Comeron, analyzes the syn-rift geometry of the northeastern part of the Neuquén Basin in Neuquén Province using seismic data and a discrete element modeling approach. They present a regional map showing the half-graben faults and transfer zones that developed during the Triassic to lower Jurassic synrift stage. They argue that these faults result from differential subsidence and that the anticlines and synclines associated with them are not due to later tectonic inversion as is the case further west in the Neuquén basin. The eighth and ninth chapters consider aspects of late Oligocene to middle Miocene arc and retroarc deformation and magmatism. The eighth chapter, by W.M. Burns, T.E. Jordan, P. Copeland, and S.A. Kelley, presents a synthesis of the Cura-Mallín basin which is one in a chain of late Oligocene to early Miocene intra-arc sedimentary basins that formed along the Andes between 33° and 43°S. Facies variations, stratal thickness patterns, structural analyses, apatite and zircon fission-track and 40Ar/ 39Ar ages are used to argue that the basin formed by normal faulting, with little or no strike-slip influence. The model is argued to apply to the entire chain of intra-arc basins. The ninth chapter, by S.M. Kay and P. Copeland, presents a synthesis of
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Preface early to middle Miocene backarc magmatism between 36°S and 37°S with new major, trace element and isotopic data, and 40Ar/ 39Ar ages. The lack of an arc chemical signature in 24 to 20 Ma alkaline magmas and its appearance in amphibole-bearing mafic lavas after 20 Ma is linked to the initial shallowing of the subducting Nazca plate. This information is put together with the space-time distribution and chemical-isotopic characteristics of Oligocene to middle Miocene retroarc magmas between 33° and 43°S to argue that a regional change from extensional to contractional deformation after 20 Ma is best attributed to an accelerated rate of westward drift of South America over the underlying mantle. The tenth chapter, by S.M. Kay, O. Mancilla, and P. Copeland, examines the evolution of the late Miocene Chachahuén volcanic complex near 37°S that is located some 500 km east of the modern Chile trench. A detailed analyses of mineralogical, geochemical and isotopic data and 40Ar/39Ar ages along with previously unpublished K/Ar ages and mapping shows that the magmatic evolution of this 7.3–4.8 Ma hornblende-bearing basaltic to dacitic nested caldera complex is best explained by formation during a latest Miocene peak in a transient shallowing of a portion of the subducting Nazca plate under the northern Neuquén Basin. The uplift of the Sierra de Chachahuén is also argued to have occurred at this time. Attention is drawn to chemical and structural parallels with similar age volcanic complexes and uplifts far to the east of the trench over the modern Chilean flatslab region between 28°S and 33°S. The cause of transient shallow subduction could be subduction of a small oceanic plateau. Chapters eleven to thirteen principally deal with the Pliocene to Holocene tectonic evolution of the Main Cordillera and the eastern slope of the Neuquén Andes. In chapter eleven, A. Folguera, V.A. Ramos, E.F. González Díaz, and R. Hermanns present five structural transects between 37°S and 37°30´S that demonstrate the existence of late Miocene to Quaternary folds and thrusts that were previously largely unrecognized on the eastern slope of the Andes. They call this region the Guañacos fold and thrust belt and demonstrate that contractional inversion of Oligocene to Miocene extensional structures has affected even Quaternary volcanic rocks in the present orogenic front that is located immediately east of the Southern Volcanic Zone arc. The inversion is mechanically linked with the La Laja strike-slip fault system in the intra-arc in Chile. In chapter twelve, A. Folguera, T. Zapata, and V.A. Ramos describe the contrasting Pliocene to Holocene deformational styles and topographic features north and south of 37.5°S on the eastern slope of the Andes between 36°S and 39°S. They use published data along with new field and seismic evidence to discuss four extensional depocenters that developed contemporaneously with the Guañacos fold-and-thrust belt. They argue that these contrasting deformational styles, which are difficult to reconcile with the nearly constant convergence parameters and slab configuration along the modern margin, can be reconciled with different rates of trench roll back in response to different late
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Miocene to Holocene changes in Wadati-Benioff zone geometry north and south of 37.5°S . The thirteenth chapter, by F. Miranda, A. Folguera, P.R. Leal, J.A. Naranjo, and A. Pesce, deals with the age of volcanism and deformation in the upper Pliocene to lower Quaternary volcanic chain east of the modern Andean front between 36.5° and 38°S. The authors present new K/Ar ages, generally bracket a Miocene deformation event between 9 and 6 Ma, and show that the volcanic centers erupted in two contrasting structural settings. One group is associated with north northwest-trending contractional faults and the other with northeast-trending extensional faults. The difference in structural styles can be explained by strain partitioning in the region. The fourteenth and fifteenth chapters deal with volcanism in the Pliocene to Quaternary Southern Volcanic Zone arc. In chapter fourteen, L.E. Lara and A. Folguera present an overview of the complex arc-backarc magmatic system that developed on the western margin of the Neuquén Basin during the late Cenozoic. They use new 40Ar/39Ar ages, compiled K/Ar ages and geochemical data, and the regional tectonic framework to argue that paired volcanic belts south of 38°S that have been used as evidence for a westwardly migrating volcanic front actually reflect arc broadening followed by narrowing. In their view, the arc front remained stationary. The narrowing of the arc is argued to correlate with a decrease in plate convergence rate and dextral transpression. The last chapter, by J.C. Varekamp, J. Maarten de Moor, M.D. Merrill, A.S. Colvin, A.R. Goss, P.Z. Vroon, and D.R. Hilton, presents a detailed look at the geochemical and isotopic evolution of the Pliocene to Holocene Caviahue-Copahue volcanic complex, which sits at the junction where the eastern Pliocene chain diverges from the main Southern Volcanic Zone arc. The paper presents a large new data set with major and trace elements and Pb, Sr, Nd, and He isotopic data. The authors argue that chemical and isotopic changes between the Pliocene Caviahue complex and the Pleistocene Copahue volcano require a change in magma source components and a change from a fluid fluxing melt regime to a sediment melting regime. They note that the nature of mantle melting could have switched from flux melting to decompressional melting with a change to an extensional regime in the region. The comprehensive picture of the structural, magmatic, sedimentological, and tectonic evolution of the Andes between 35°S and 39°S that emerges through these chapters provides a new framework and perspective on the evolution of the distinctive Andean-type accretionary orogen in this region. The conclusions have major implications for the Cenozoic evolution of the Neuquén foreland basin. ACKNOWLEDGMENTS The editors would like to sincerely thank the reviewers for their constructive reviews and comments on the papers. In alphabetical order, the reviewers were: Adriana Bermúdez (Universidad Nacional del Comahue, Argentina), Richard All-
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mendinger (Cornell University, USA), Susan Beck (University of Arizona, USA), Ben Brooks (University of Hawaii, USA), Mathew Burns (U.S. Geological Survey, USA), Carlos Costa (Universidad Nacional de San Luis, Argentina), Peter Cobbold (University of Rennes, France), Beatriz Coira (Universidad Nacional de Jujuy, Argentina), John Davidson (Durham University, UK), Gloria Eisenstadt (University of Texas at Arlington, USA), Luis Fauque (SEGEMAR, Argentina), Todd Feeley (Montana State University, USA), Estanislao Godoy (SERNAGEOMIN, Chile), Brian Horton (University of California at Los Angeles, USA), Robert Kay (Cornell University, USA), Estanislao Kozlowski (Pan American Oil, Argentina), Eduardo
Llambías (Universidad de La Plata, Argentina), Leopoldo López Escobar (Universidad de Chile, Concepción, Chile), Andrew Meigs (Oregon State University, USA), Constantino Mpodozis (Sipetrol, Chile), Jorge Muñoz (SERNAGEOMIN, Chile), Eric Sandvol (University of Missouri, USA), Charles Stern (University of Colorado, USA), Anthony Tankard (Consulting Geologist, Canada), Stuart Thompson (Yale University, USA), Gustavo Vergani (Repsol-YPF, Argentina), and Jan Witte (Wintershall Energía, SA, Argentina). Suzanne Mahlburg Kay Víctor A. Ramos
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Geological Society of America Special Papers Overview of the tectonic evolution of the southern Central Andes of Mendoza and Neuquén (35° −39°S latitude) Victor A. Ramos and Suzanne Mahlburg Kay Geological Society of America Special Papers 2006;407;1-17 doi: 10.1130/2006.2407(01)
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Geological Society of America Special Paper 407 2006
Overview of the tectonic evolution of the southern Central Andes of Mendoza and Neuquén (35°–39°S latitude) Víctor A. Ramos* Laboratorio de Tectónica Andina, Universidad de Buenos Aires, Consejo Nacional de Investigaciones Científicas y Técnicas, Buenos Aires, Argentina Suzanne Mahlburg Kay* Institute for the Study of the Continents and Department of Earth and Atmospheric Sciences, Snee Hall, Cornell University, Ithaca, New York 14853, USA
ABSTRACT The southern Central Andes of Argentina between 35° and 39°S latitude can be divided into two sectors with contrasting geological histories. The boundary between the sectors coincides with the Cortaderas lineament. North of the Cortaderas lineament, the Andes record a foreland expansion of arc magmatism that is associated with contractional deformation in the Malargüe fold-and-thrust belt, and subsidence of the Río Grande foreland basin between 15 and 5 Ma. The peak expansion of deformation into the foreland occurred as late Miocene magmatic arc rocks erupted more than 500 km east of the trench and the San Rafael basement block was uplifted in central Mendoza. This stage was followed by the collapse of the basement uplift by normal faulting, the retreat of the magmatic arc, and the eruption of widespread late Pliocene to early Pleistocene within-plate lava flows in the Payenia region. Extensive Quaternary calderas and rhyolitic domes along the axis of the main Andes reflect crustal melting associated with basaltic underplating. In contrast, the structural evolution of the sector south of the Cortaderas lineament is dominated by the Late Cretaceous development of the Agrio fold-and-thrust belt, which underwent minor reactivations in the Eocene and the late Miocene. The post-Miocene Guañacos fold-and-thrust belt that has since developed along the axis of the main Andes concentrates neotectonic contraction. Arc magmatism in this sector is largely restricted to the axial area of the Andes. Both the sectors north and south of the Cortaderas lineament show evidence of an important episode of extension during the Oligocene to early Miocene, and for renewed extension in the Pliocene and the Pleistocene. The contrasting geological histories north and south of the Cortaderas lineament reflect differences in the geometry of the subducting plate, variations in crustal rheologies inherited from a more restricted distribution of Mesozoic rifts in the northern than the southern segment, and variations in the trench roll-back velocity through time. Keywords: Andes, Neuquén, Malargüe, Agrio, Payenia, flat-subduction, extension. *E-mails:
[email protected];
[email protected]. Ramos, V.A., and Kay, S.M., 2006, Overview of the tectonic evolution of the southern Central Andes of Mendoza and Neuquén (35°–39°S latitude), 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 latitude): Geological Society of America Special Paper 407, p. 1–17, doi: 10.1130/2006.2407(01). For permission to copy, contact
[email protected]. ©2006 Geological Society of America. All rights reserved.
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INTRODUCTION The various papers in this volume show that the segment of the Andean Cordillera between 35° and 39°S has distinctive characteristics when compared with other parts of the ArgentineChilean Andes. Most of the Andean chain is under compression, and the present orogenic front is located between a retroarc fold-and-thrust belt and an undeformed foreland region. Magmatic rocks are generally concentrated along an active magmatic arc, except in areas of shallow subduction where volcanism is absent (e.g., Jordan et al., 1983; Kay et al., 1988). In contrast, the segment discussed here (Fig. 1) has some peculiarities, including: (1) a large amount of late Cenozoic magmatic activity in the foreland region (Bermúdez et al., 1993); (2) an active thrust front located west of an inactive Late Cretaceous to Miocene fold-and-thrust belt (Folguera et al., 2004; Ramos and Folguera, 2005); and (3) widespread Pliocene and Pleistocene extension (Polanski, 1963; González Díaz, 1964; Kozlowski et al., 1993; Folguera and Ramos, 2000, 2001, 2002). The objective of this contribution is to present a synthesis of the tectonic evolution of the Neuquén and southern Mendoza segments of the Andes between 35° and 39°S based on information in recent publications and in the chapters in this volume. This synthesis shows that geologic observations in the northern part of the region fit with steepening of the subduction zone fol-
lowing a short period of flat-slab subduction during the Miocene as proposed by Kay (2001a, 2001b, 2002; Kay et al., this volume, chapter 2) and that those in the southern part fit with an oscillatory behavior of the subduction zone. MAIN GEOLOGICAL FEATURES The Andes between 35° and 39°S can be divided into two regions with distinctive evolutionary characteristics. These regions are bounded by a northwest-trending structural feature that was first described by Groeber (1938) and is referred to as the Cortaderas lineament (Ramos, 1981; Ramos and Barbieri, 1989). The Cortaderas lineament is defined by northwest-trending faults that can be traced to the Southern Volcanic Zone arc front where they control the recent craters of the Chillán volcano. The lineament was interpreted as a basement boundary that was reactivated during Andean deformation by Ramos (1981). Other authors consider the Cortaderas lineament to be a reactivated Paleogene north-verging thrust system (Cobbold and Rossello, 2003). Another important observation associated with the Cortaderas lineament is the concentration of Cenozoic magmatic activity to the north and the near absence of Cenozoic volcanism to the south (Ramos and Barbieri, 1989). Kay (2001a, 2001b; Kay et al., this volume, chapter 2) argued that the Cortaderas lineament marks the southern limit of a Miocene shallow subduction zone.
Figure 1. Map from the trench to the backarc showing the main geological provinces and structural features of the south-central Andes between 34°and 40°S. Differences in the regions north and south of the Cortaderas lineament are discussed in the text. Position of the Cortaderas lineament is based on Ramos and Barbieri (1989). FTB—fold-and-thrust belt.
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Tectonic evolution of the southern Central Andes, Mendoza and Neuquén The major geologic features of the regions north and south of the Cortaderas lineament are illustrated in the figures and summarized herein. Region North of the Cortaderas Lineament The major geologic features in the region between 37°S and 34°S, to the north of the Cortaderas lineament, are illustrated in Figures 2 and 3. The general characteristics of the area
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just north of the Cortaderas lineament are described in this volume by Kay and Copeland (chapter 9), Kay et al. (chapters 2 and 10), and Folguera et al. (chapters 11 and 12). The discussion herein integrates these features with those farther north in the province of Mendoza. From west to east, the main morphotectonic elements of the region include the main Andean range (Principal Cordillera), the Río Grande foreland basin, the San Rafael basement block, and the widespread Pliocene to Quaternary Payenia backarc volcanic field.
Figure 2. Generalized geological map of parts of the provinces of Mendoza, Neuquén, and La Pampa in Argentina showing the Miocene geologic features of the region north of the Cortaderas lineament discussed in the text. Post-Miocene structures and volcanic rocks in the foreland are shown in Figure 3. Bold dashed lines are principal highways.
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Figure 3. Map of the same region as in Figure 2 showing the principal Pliocene and younger structures and volcanic rocks. Lines show locations of cross sections shown in Figures 4 and 5. Short dashed lines are boundaries between the Argentine provinces of Mendoza, Neuquén, and La Pampa. Longer dashed lines are principal highways (symbol is for Highway 40).
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Tectonic evolution of the southern Central Andes, Mendoza and Neuquén Principal Cordillera The Principal Cordillera straddles the Chilean-Argentine border zone (Figs. 1 and 2) and includes the active Andean Southern Volcanic Zone arc front, which at these latitudes (35°–39°S) is located on the Chilean side of the range. Generally, the Southern Volcanic Zone has been divided into three segments based on petrologic and geologic studies (e.g., López Escobar, 1984; Hildreth and Moorbath, 1988; Dungan et al., 2001). These segments are known as the northern Southern Volcanic Zone (Tupungato to Maipo, 33.5°–34.5°S latitude), the transitional Southern Volcanic Zone (Palomo to Tatara–San Pedro, 35°–36°S), and the southern Southern Volcanic Zone (Longaví to Hudson, 36°–46°S). Among other differences, crustal thicknesses are considered to decrease from 65–60 km in the northern Southern Volcanic Zone to ~42–35 km in the southern Southern Volcanic Zone (e.g., Hildreth and Moorbath 1988; Ramos et al., 2004). Recent studies have demonstrated a complex history of forearc subduction erosion and consequent late Cenozoic migration of the magmatic arc toward the foreland in the segment north of 36°S (Kay et al., 2005). North of the Cortaderas lineament, the eastern slope of the Principal Cordillera coincides with the Las Loicas trough (named by Folguera et al., this volume, chapter 12). The Las Loicas trough is a Pliocene to Quaternary volcano-tectonic basin located east of the active Southern Volcanic Zone arc that is controlled by extensional north-northwest–trending faults and filled with thick sequences of Quaternary ignimbrites, lavas, and ashfall deposits derived from large silicic volcanic centers like the Bobadilla, Varvarco, and Domuyo calderas (Fig. 3; see Hildreth et al., 1999; Folguera et al., this volume, chapter 12). East of the Las Loicas trough lies the thick-skinned Malargüe fold-and-thrust belt (Fig. 2), which deforms Mesozoic marine and continental sequences (e.g., Kozlowski et al., 1993).
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The structure of this belt is dominated by large double-plunging anticlines that are cored by the Choiyoi Group (Permian-Triassic) volcanic and plutonic rocks, which constitute the exposed basement of the region. The Mesozoic strata in the belt are separated by an angular unconformity from pre–late Miocene Tertiary volcanics and granitoids, which are in turn unconformably overlain by thick sequences of late Miocene to Pliocene volcanics (Gerth, 1931; Ramos and Nullo, 1993). Tertiary synorogenic conglomerates and sandstones occur in the Principal Cordillera as remnants of the western part of a foreland basin. Cobbold and Rossello (2003) discussed the evolution of the Paleogene deposits. The volume of the Paleogene deposits is relatively minor compared to that of Neogene sequences that are widely preserved in the eastern foothills of the Principal Cordillera from the Río Diamante valley in the north to the Río Grande valley in the south (Fig. 2). The Neogene sequences are interpreted as synorogenic deposits associated with the Miocene uplift and shortening of the main Andean range. According to Kraemer et al. (2000), an older sequence that is constrained between 15.1 Ma at the base and 6.7 Ma at the top by K/Ar ages on interbedded volcanic layers is separated by an angular unconformity from a younger sequence, the age of which, at the base, is constrained to be between 6.7 Ma and 5.4 Ma. Undeformed late Pliocene to Quaternary sequences, which include andesitic and basaltic lava flows, unconformably overlie the Miocene deposits. Neogene normal faults are found along the eastern foothills of the Principal Cordillera. One of these is the Infiernillo fault (Kozlowski et al., 1993; Dajczgewand, 2002) that intersects the Río Salado valley (Figs. 3 and 4). As shown in Figure 4, the Infiernillo fault is a west-dipping normal fault that separates late Miocene deposits from Mesozoic sedimentary strata. Evidence for activity on this fault in Quaternary times comes from distinct pulses of basaltic lavas that have erupted along the fault (Fig. 4).
Figure 4. Cross section across the El Infiernillo normal fault along the Río Salado valley (modified from Dajczgewand, 2002). Location of section is shown in Figure 3.
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Río Grande Foreland Basin East of the Principal Cordillera, synorogenic deposits have accumulated in the large Río Grande foreland basin, which extends from the Río Diamante valley in the north to the southern limit of Mendoza province (Fig. 2). The basin fill includes more than 2000 m of synorogenic deposits that accumulated in two distinct depocenters located north and south of the Río Atuel (Yrigoyen, 1993). The Castillos de Pincheira outcrops, a few kilometers west of the city of Malargüe, expose both Paleogene and Neogene synorogenic sequences. Facies analyses of the Paleogene unit indicate a western provenance. Facies analyses of the Miocene deposits show that they constitute a coarseningupward sequence with both an eastern and western provenance. The thick coarse deposits at the top of the sequence have a dominantly eastern provenance (Kraemer and Zulliger, 1994) and contain basement clasts that record the uplift of the San Rafael block (Fig. 2, see following). These sequences are unconformably covered by thick late Pliocene to Quaternary deposits. Subsidence associated with extension has affected the region of the Río Grande basin in recent times. Active subsidence is presently occurring in the Laguna Llancanelo depression (Fig. 3) on the eastern side of the basin. This depression is bounded on the east by a normal fault (Manceda in Kozlowski et al., 1993). To the north is the Alto Tunuyán depression, which is interpreted as a half-graben (Fig. 3) that is filled by Quaternary clastic sediments. As recognized by Polanski (1963), the eastern side of the Alto Tunuyán depression is bounded by north-south–trending normal faults that affect Pliocene deposits and that can have throws of more than 700 m. One of these is the Cerro Negro de Capiz fault that affects late Pliocene deposits (Yrigoyen, 1993). The Extenso del Campo Bajo valley occurs in the region between the Alto Tunuyán depression and the Laguna Llancanelo.
San Rafael Basement Block Further east is the San Rafael basement block (Fig. 2). This block consists of Middle Proterozoic metamorphic rocks and tightly folded Paleozoic deposits (Moreno Peral and Salvarredi, 1984) that are cut and unconformably overlain by Permian to Triassic granitoids and volcanic rocks of the Choiyoi Group (e.g., Kay et al., 1989). The presence of an old peneplain carved on these basement rocks was first noted by Polanski (1954). The time of uplift and exposure of the erosional surface is constrained by synorogenic deposits of the Río Grande foreland basin that form a bajada of low-energy clastic deposits, which were derived from the Andean foothills. The time of deposition of the sediments is constrained by the presence of mammalian fossils (Soria, 1984) with a middle Miocene Colloncurense age (15–12 Ma; Pascual et al., 2002) and the K/Ar ages of 15.1 Ma and 6.7 Ma in the tuffs in the sequences dated by Kraemer et al. (2000). A rapid period of uplift at 5 Ma is indicated by Pliocene synorogenic deposits with an eastern provenance from the Río Grande basin. This uplifted peneplain was subsequently cut by Pliocene and Quaternary normal faults described by Narciso et al. (2001). Among these normal faults is the west-dipping El Carrizalito fault (Figs. 3 and 5), which indicates activity in the late Pliocene to early Pleistocene recognized by González Díaz (1964). Another is the Llancanelo fault that bounds the western margin of the block further to the south. The eastern margin of the San Rafael block is still actively shortening, as indicated by recent seismic activity as exemplified by the large Malvinas earthquake in the vicinity in 1929 (Bastías et al., 1993). The thrust front, which does not show a surface rupture, is expressed by a frontal monocline developed in Pleistocene basalts (Costa et al., 2004).
Figure 5. Cross section across the San Rafael block showing the Carrizalito fault and the late Miocene peneplain, which is broken by normal faults. One of these faults controls the position of the late Pleistocene basalts that erupted from the Cerro Negro volcano. The late Miocene thrust on the eastern side is inactive. Location of section is shown in Figure 3.
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Tectonic evolution of the southern Central Andes, Mendoza and Neuquén Retroarc Volcanic Field The large retroarc volcanic province named Payenia by Polanski (1954) covers vast regions of the foreland between 35° and 37°S (Fig. 3). The maximum volumes of volcanic rocks erupted in the region between the Volcán Nevado and the Payún Matru caldera. These volcanic rocks pinch out to the north where they partially cover the San Rafael block. Regional studies by González Díaz (1972a, 1972b) and Bermúdez et al. (1993) indicate that these late Miocene to Quaternary volcanic rocks erupted in two distinct volcanic episodes. The first episode produced the hornblende-bearing mafic andesites to rhyodacites found in the Cerro Plateado (35°40′S) and Sierra de Chachahuén (37°05′S) volcanic fields, which are located some 500 km east of the modern trench (Fig. 2). The andesitic to rhyolitic volcanic rocks in the Cerro Plateado field are considered to be late Miocene in age (Bermúdez, 1991). Volcanic rocks in the Sierra de Chachahuén field farther south have an arc-like magmatic chemistry and range in age from ca. 7.2 to 4.8 Ma (Kay, 2001a, 2001b; Kay et al., this volume, chapter 9). Younger volcanic rocks from the Volcán Nevado (3810 m) in the region of the Plateado field (Fig. 3) are dominantly composed of trachyandesite and are considered to be Pliocene in age (Bermúdez, 1991). The second episode produced the extensive late Pliocene to Quaternary Payenia volcanic field that is temporally associated with the Cerro Diamante (2354 m), Cerro Payún (3680 m), Payún Matru caldera, and Auca Mahuida volcanic complexes (Figs. 3 and 6; see Bermúdez et al. 1993). Basaltic flows in these fields can reach lengths of over 200 km. The volcanic rocks of this episode generally have an alkaline within-plate chemical signature and are considered to be associated with an extensional regime (Muñoz et al., 1989; Bermúdez et al., 1993; Kay, 2001a, m2001b; Kay et al., this volume, chapter 2). Region South of the Cortaderas Lineament The major geologic features of the region south of the Cortaderas Lineament are illustrated in Figures 1 and 6. In striking contrast with the region north of the Cortaderas lineament, the retroarc east of the Loncopué trough is essentially devoid of Neogene volcanic rocks and synorogenic deposits (Fig. 6). Many of the general characteristics of this region are described in the papers in this volume by Mosquera and Ramos (chapter 5), Zamora Valcarce et al. (chapter 6), and Folguera et al. (chapters 11 and 12). From west to east, the main morphotectonic elements of the region are the Principal Cordillera, the Loncopué trough, the Agrio fold-and-thrust belt, the Chihuidos high, the Añelo basin, and the Neuquén Embayment. Principal Cordillera The Principal Cordillera south of the Cortaderas lineament is located within the southern segment of the Southern Volcanic Zone arc. Quaternary to Holocene arc volcanic complexes in this region include the centers at Chillán, Antuco, Copahue,
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Lonquimay, and Llaima (Fig. 6), which were assigned to the Lonquimay volcanic chain by Burckhardt (1900). These centers overlie Mesozoic marine sediments, Cretaceous plutonic rocks, and Oligocene to Miocene sedimentary and volcanic rocks that are unconformably covered by early Pliocene Cola de Zorro Formation volcanic rocks (e.g., Vergara and Muñoz, 1982; Burns et al., this volume, chapter 8). The Lonquimay chain is flanked to the east by the series of late Pliocene to early Quaternary volcanoes that Burckhardt (1900) assigned to the Pino Hachado chain. Most of the eastern flank of the Principal Cordillera north of the Copahue volcano is affected by the contractional deformation that created the late Miocene to Quaternary Guañacos fold-and-thrust belt (Folguera et al., 2004, this volume, chapter 11). Loncopué Trough The Loncopué trough, which is immediately east of the Principal Cordillera (Fig. 6), is an extensional basin associated with a thick cover of Pleistocene basaltic lavas and late Pleistocene– Holocene monogenic volcanic cones (Ramos, 1978; García Morabito, 2005; Folguera et al., this volume, chapter 12). The volcanic rocks have subdued-arc geochemical signatures (Muñoz and Stern, 1988). The Loncopué trough coincides with an important region of crustal attenuation that is discussed by Yuan et al. (this volume, chapter 3). Agrio Fold-and-Thrust Belt The next major feature to the east is the Agrio fold-andthrust belt (Figs. 1 and 6), which was divided into two sectors by Ramos (1998). The western sector is composed of Jurassic and Early Cretaceous marine deposits that were deformed in a thick-skinned belt, which was produced by the inversion of Early Jurassic half-grabens (Vergani et al., 1995; Folguera et al., 2002) during the Late Cretaceous. These sequences are intruded and overlain by Late Cretaceous to Eocene magmatic rocks (Franchini et al., 2003; Zamora Valcarce et al., this volume, chapter 6). The eastern sector of the Agrio belt consists of Early Cretaceous marine sedimentary rocks that were deformed in a Late Cretaceous thin-skinned belt that detached in Late Jurassic evaporite deposits. Both parts of the Agrio belt were tectonically reactivated and last deformed in the middle to late Miocene (Zapata and Folguera, 2005). Chihuidos High Farther east, the Chihuidos high is a basement arch that is largely covered by Late Cretaceous red bed sequences, which are synorogenic foreland basin deposits associated with the Agrio fold-and-thrust belt. The age of the lower part of the synorogenic sequence is constrained by zircon fission-track ages of 88 ± 3.9 Ma in the Huincul Formation on the southern margin of the basin at Cerro Policía in the province of Río Negro (Corbella et al., 2004). These ages imply that deformation in the Agrio fold-and-thrust belt had started by the Late Cretaceous (before Turonian-Santonian times). According to Mosquera and
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Figure 6. Generalized geological map of the northern part of Neuquén province in Argentina showing the geologic features along and south of the Cortaderas lineament discussed in the text.
Ramos (this volume, chapter 5), the presence of deformed middle Miocene synorogenic deposits on the Chihuidos high shows that the block was uplifted by tectonic inversion of normal faults during the late Miocene. Small post-tectonic basaltic cones with ages of 4.5 ± 0.5 Ma (Ramos and Barbieri, 1989) and an intraplate alkaline chemistry (Kay et al., 2004) occur near the northern boundary of the Chihuidos high (Fig. 6). Añelo Basin The sediments in the Añelo basin (Fig. 6) to the east record the Miocene uplift of the Chihuidos high. This basin represents a foredeep with a fill of a few hundred meters of late Miocene to Pliocene synorogenic deposits (Mosquera and Ramos, this volume, chapter 5), which include the strata exposed at Barranca del Palo and in the Sierra Blanca (Uliana, 1978). Most of the synorogenic sediments that formed in the Miocene
bypassed the Añelo basin and Neuquén Embayment and ended up as far east as the continental margin (Folguera et al., 2005). Neuquén Embayment East of the Agrio fold-and-thrust belt and south of the Añelo basin, Mesozoic sedimentary rocks associated with the Neuquén Embayment are preserved (Fig. 6). These sequences include a large expanse of marine sediments that were deposited and subsequently covered by Late Cretaceous synorogenic deposits. Thin sequences of continental and shallow marine strata of Maastrichtian to Paleogene age represent the first transgression derived from the Atlantic Ocean after the Early Cretaceous opening of the South Atlantic. The more northern part of the embayment is partially covered by Neogene alkaline basaltic lavas associated with volcanic centers north of the Cortaderas lineament.
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Tectonic evolution of the southern Central Andes, Mendoza and Neuquén MAGMATIC AND DEFORMATIONAL HISTORY Important differences in the regions north and south of the Cortaderas lineament are evident in the sequential episodes of contractional and extensional deformational events that affected this region of the Andes. These differences are discussed in the following within the general context of the tectonic cycles first proposed by Groeber (1953, and references therein). Early Jurassic–Early Cretaceous Extensional Stage Most of the Early Jurassic to Early Cretaceous Andean margin at these latitudes was associated with negative trench roll-back velocity of the subduction zone (see Mpodozis and Ramos, 1989; Ramos, 1999). Normal faults were dominant along the main Andean chain, and volcanic rocks were widespread on both slopes of the Main Andean Cordillera. During this time, deep marine depocenters accumulated thick deposits that interfingered with basaltic and andesitic rocks erupted in an extensional-type magmatic arc. The main difference between the regions north and south of the Cortaderas lineament is that rift systems are more aerially restricted to the north. In detail, widespread grabens and halfgrabens found in the foreland in the south are abruptly replaced by a much narrower zone of rift-related structures in the north. The faulting in these rifts, which started in the Triassic and continued into the Early Jurassic, preceded the inception of subduction along the Pacific margin. The wider zone of rifts in the south coincides with a distinctive basement fabric. An important feature of this basement fabric is the N70°E-trending Huincul fault system that corresponds with the Huincul Ridge (Figs. 1 and 6) and marks a major crustal boundary. Mosquera and Ramos (this volume, chapter 5) interpret this boundary as a Paleozoic suture marking the collision between an allochthonous Patagonian terrane to the south and the Gondwana continent to the north. Evidence that the Huincul Ridge parallels the suture comes from aeromagnetic (Chernicoff and Zappettini, 2004), gravity (Kostadinoff et al., 2005), and seismic (Mosquera and Ramos, this volume, chapter 5) data. Late Cretaceous Contractional Stage Important changes in magmatism and deformation occurred near the end of the Early Cretaceous. One of these changes was the shutdown of the elongated magmatic belt that produced the 93–72 Ma granitoids exposed along the Chilean side of the Principal Cordillera (see Ramos and Folguera, 2005). South of the Cortaderas lineament, the end of this magmatic stage coincided with the eastward migration of the Late Cretaceous magmatic arc and its reestablishment in the western part of the Agrio fold-and-thrust belt. A contemporaneous migration of the frontal arc front did not occur in the northern sector. Another major development at this time was that the deformational
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regime became contractional throughout the region, with the most intense deformation occurring in the southern sector. The more extensive previous rifting in the southern region likely left an extended broken basement that was more susceptible to deformation. Support for major Late Cretaceous contractional deformation south of the Cortaderas lineament comes from several lines of reasoning. First, 40Ar/39Ar ages obtained by Zamora Valcarce et al. (this volume, chapter 6) on pyroclastic breccias and andesitic lavas unconformably overlying deformed Early Cretaceous marine rocks show that deformation had begun by 77 Ma in the western Agrio fold-and-thrust belt. Second, 40Ar/39Ar ages obtained by Zamora Valcarce et al. (this volume, chapter 6) on east-west–trending andesitic dikes suggest that compression along the western margin of the Neuquén Embayment had begun by 100 Ma. Third, zircon fission-track ages in Neuquén Group strata reported by Corbella et al. (2004) show that foreland basin sedimentation had begun by 88 Ma. Finally, fossils in Neuquén Group rocks indicate the development of a deep, eastward-thinning, Late Cretaceous foreland basin related to contractional deformation in the Agrio fold-and-thrust belt. Evidence for Late Cretaceous contractional deformation north of the Cortaderas lineament comes from the observation that andesitic volcanic rocks south of Cerro Domuyo and east of the Cordillera del Viento (Fig. 6) with K/Ar ages of 71.5 ± 5 Ma postdate deformed Cretaceous deposits (Llambías et al., 1979). Other evidence comes from a 69.09 ± 0.13 Ma 40Ar/39Ar biotite cooling age from the Varvarco pluton, which Kay (2001b) and Kay et al. (this volume, chapter 2) interpret as an uplift age for the Cordillera del Viento. Burns (2002) argued from fissiontrack data on zircons that the uplift of the Cordillera del Viento had begun by 70 Ma, and possibly before 80 Ma. Paleogene Compressional Stage There are still uncertainties concerning the distribution of magmatic rocks and the extent of contractional deformation that took place during the Paleogene. Well-preserved Eocene volcanic rocks occur along the Chilean Central Valley in the northern sector, but as noted by López-Escobar and Vergara (1997), similar-age volcanic rocks have not been found in Chile between 36° and 37°30′S. Paleogene volcanic rocks are present at these latitudes in Argentina (Franchini et al., 2003; Ramos and Folguera, 2005). Kay et al. (this volume, chapter 2) argue for a small eastward shift of the volcanic front in the Paleogene based on the age, chemistry, and distribution of volcanic rocks in northern Neuquén. The scarcity of Paleogene synorogenic deposits across the region is linked to the question of the importance of Paleogene contractional deformation. Volumes of these deposits are low north of the Cortaderas lineament and even more limited to the south. A notable exception is in the Pampa del Agua Amarga region in the southern Agrio fold-and-thrust belt, where late Paleocene to early Eocene pyroclastic and tuffaceous deposits
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in the Puesto Burgos Formation unconformably overlie sediments of the Late Cretaceous Neuquén Group (Leanza and Hugo, 2001). The age of these tuffs is constrained by plant fossils. Overall, it is apparent that large parts of the Agrio fold-andthrust belt and the adjacent Chihuidos high were positive areas during the Paleogene and that some synorogenic sediments accumulated in a foreland basin in the north. Cobbold et al. (1999) and Cobbold and Rossello (2003) appealed to transpression to reconcile the relatively small amounts of synorogenic deposits found for this time with the important episode of Eocene contractional deformation that they proposed. Their argument for Eocene deformation was largely based on the orientation of bitumen veins in the Agrio fold-and-thrust belt east of the Cordillera del Viento. Farther south, evidence for Eocene uplift along the Cordillera Principal south of 38°S comes from apatite fission-track ages of 40.6 ± 4.5 Ma (Gräfe et al., 2002). Oligocene to Early Miocene Extensional Stage During the Oligocene to early Miocene, the Neuquén Andes were characterized by generalized extension in the forearc (e.g., Cisternas and Frutos, 1994; Stern et al., 2000), arc (e.g., Vergara et al., 1997; Burns et al., this volume, chapter 8), and retroarc (e.g., Folguera et al., 2003). The eruption of early Miocene alkaline basaltic volcanic rocks at distances up to 500 km from the modern trench has been used as evidence that crustal attenuation in an extensional regime extended far into the foreland (Ramos and Barbieri, 1989; Kay, 2001b; Kay and Copeland, this volume, chapter 9). This period has been associated with a period of negative trench roll-back by Kay and Copeland (this volume). Middle to Late Miocene Contractional Stage The middle to late Miocene was a time of contractional deformation across the region. Important differences occur north and south of the Cortaderas lineament. To the north, magmatism was widespread, the Malargüe fold-and-thrust belt propagated eastward, and synorogenic deposits were widely distributed. To the south, magmatism was confined to the region near the arc axis, and contractional deformation was restricted to inversion of the Cura Mallín basin in the Principal Cordillera and to minor reactivation of the Agrio fold-andthrust belt. The most dramatic changes occurred in the northern segment. Beginning in the region of the Principal Cordillera, the eastward broadening of the magmatic arc at the latitude of central and southern Mendoza was initially recognized by Gerth (1931). This eastward expansion of magmatism occurred at the time of important crustal shortening and uplift in the Malargüe fold-and-thrust belt, and the subsidence that led to the accumulation of the thick Neogene synorogenic deposits in the Río Grande basin. The eastward propagating sequence culminated
in the contractional uplift of the San Rafael block by the end of the Miocene and the eruption of the Cerro Plateado and associated volcanic centers far to the east of the trench (Delpino and Bermúdez, 1985). Bermúdez (1991) pointed out that the late Miocene arc in central Mendoza was over 200 km wide. The eastward expansion of the Miocene volcanic arc across southernmost Mendoza and northern Neuquén that culminated in the eruption and uplift of the Sierra de Chachahuén between 7.8 and 4.8 Ma is discussed by Kay (2001a, 2001b) and Kay et al. (this volume, chapters 2 and 10). The presence of sparse synorogenic deposits in the foreland south of the Cortaderas lineament is interpreted as indicating renewed contraction in this region. Among these deposits are those at Rincón Bayo that unconformably overlie Paleogene sediments in the Chihuidos high (Zapata et al., 2003) and conglomerates and sandstones containing middle Miocene mammal fossils (Repol et al., 2002). Based on the small volume of Miocene compared to Cretaceous synorogenic deposits, Miocene contraction in this region is considered to have been less important than in the Cretaceous (Ramos, 1998; Mosquera and Ramos, 2005, this volume, chapter 5). Pliocene Extensional Stage Differences across the northern and southern sectors are again striking in the Pliocene. The northern sector is characterized by important retroarc basaltic magmatism with a subdued arc to intraplate chemical character (Muñoz et al., 1989; Bermúdez et al., 1993; Kay, 2001b; Kay et al., 2004, and this volume, chapter 2). Large volumes of alkaline magmas erupted from fissures and important volcanic centers like the late Pliocene Payún Matru caldera. Magmatic activity was accompanied by generalized extension in the retroarc and adjacent areas, and the collapse of previously uplifted foreland areas like the San Rafael block (González Díaz, 1964; Bermúdez et al., 1993). Extension propagated from the foreland to the foothills of the Principal Cordillera as retroarc volcanoes, such as those in the Tromen region, erupted between the Principal Cordillera and the eastern foreland (Kay, 2001b; Kay et al., this volume, chapter 2). South of the Cortaderas lineament, extension was limited to the arc and western retroarc (Folguera et al., 2003, and this volume, chapter 11). The magmatic arc expanded toward the retroarc with stratovolcanoes erupting along the Pino Hachado chain (e.g., Muñoz and Stern, 1988; Lara and Folguera, this volume, chapter 14). Basaltic volcanism occurred only as far east as the Loncopué trough (Fig. 6). Pleistocene to Holocene Stage During the Pleistocene, extension propagated to the main axis of the Principal Cordillera. Large volumes of rhyolite erupted from a series of calderas and volcanic domes north of the Cortaderas lineament, including the Planchón, Calabozos, Bobadilla, Varvarco, Domuyito, and Domuyo centers (Fig. 3).
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Tectonic evolution of the southern Central Andes, Mendoza and Neuquén The rhyolitic magmas are considered to have been generated by crustal melting associated with basaltic underplating (Hildreth et al., 1999). The volcanic and pyroclastic deposits produced by these eruptions filled the fault-bounded Las Loicas trough (Fig. 3; see Folguera et al., this volume, chapter 12). Volcanic activity also took place in the retroarc in stratovolcanoes like Tromen and in alkaline monogenic basaltic and alkaline complexes like Payún Matrú and Auca Mahuida (see Holmberg, 1964; Llambías, 1966; González Díaz, 1972b; Bermúdez et al., 1993; Kay, 2001b; Kay et al., 2004, and this volume, chapter 2). During the same period, contractional deformation in the Guañacos fold-and-thrust belt straddled the western part of the Cortaderas lineament in the Principal Cordillera (Fig. 6) as the Pliocene-Pleistocene volcanic arc was thrust over Pleistocene bajadas (Folguera et al., this volume, chapter 11). Contractional deformation also occurred along the eastern margin of the San Rafael block, where thrust faults involve Quaternary lavas and the large 1929 Las Malvinas earthquake is interpreted to have had a compressional mechanism (Costa et al., 2004). Farther south, the main Pleistocene to Holocene arc volcanic activity was largely focused along the Lonquimay chain on the western slope of the Andes, where active volcanism continues in the southern sector of the Southern Volcanic Zone today. Mafic volcanism and normal faulting was largely restricted to the Loncopué trough, and evidence for Quaternary shortening is absent (see Folguera et al., this volume, chapter 12) TECTONIC EVOLUTION OF THE ANDES NORTH AND SOUTH OF THE CORTADERAS LINEAMENT The geologic history and features discussed in the previous sections are utilized in the following to propose a model for the contrasting evolution of the regions north and south of the Cortaderas lineament. Important factors in explaining the differences between these sectors are discrete crustal rheologies inherited from Paleozoic and Mesozoic events and differences in underlying subducting plate geometry. Conceptual cartoons in Figure 7, not drawn to scale, emphasize the difference in the various stages of this evolution. A more detailed discussion of the Miocene development of the region just north of the Cortaderas lineament is presented by Kay (2001a, 2001b, 2002) and Kay et al. (this volume, chapters 2 and 10). Early Rifting (Triassic to Early Jurassic) The Mesozoic evolution of the region is shown to begin in Figure 7A with the generalized extension and rifting related to the early breakup of the Pangean supercontinent. This rifting was concentrated north of the Huincul Ridge (Fig. 1), which is interpreted by Mosquera and Ramos (this volume, chapter 5) to parallel the late Paleozoic suture between the allochthonous Patagonia terrane and Gondwana. In accord with this interpretation, Mosquera and Ramos (this volume) suggest that the
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widespread extension under the Neuquén Embayment correlates with the hanging wall of the suture. This structural pattern then controls the later geometry of the Neuquén Embayment (Franzese and Spalletti, 2001). Early Subduction (Jurassic to Early Cretaceous) The initiation of subduction in the lower Jurassic at ca. 180 Ma is based on the appearance of arc magmatic rocks in Chile (see Mpodozis and Ramos, 1989; Parada 1990). The main difference between the northern and southern sectors is that arc volcanic rocks are either more abundant or are better preserved in the north (see Ramos and Folguera, 2005). The rift structures shown in the cartoon in Figure 7B are very similar to those shown in Figure 7A, since it is difficult to separate the faults related to early rifting from those produced by extension in the backarc after subduction began. Tectonic Inversion (Late Cretaceous to Paleocene) The cartoons in Figure 7C show the major differences between the southern and northern regions during the Late Cretaceous to Paleocene. In the southern region, the magmatic front migrated eastward and major contractional deformation occurred in the Agrio fold-and-thrust belt in the foreland. To the north, the arc remained stationary, and contractional deformation in the foreland Malargüe fold-and-thrust belt was less intense. In Figure 7C, the more pronounced eastward migration of the magmatic arc and the greater amount of crustal shortening in the south are tentatively linked to a relative shallowing of the Benioff zone in the south. Furthermore, as contractional deformation in both regions is primarily linked to tectonic inversion of basement faults and only secondarily to thinskinned thrusting, the more intense deformation in the south can be correlated with a greater amount of previous extension. Steepening of Benioff Zone in the Southern Sector (Oligocene to Early Miocene) Major changes occurred in the latest Oligocene to early Miocene as the Nazca plate replaced the Farallón plate, the relative convergence rate increased and became nearly normal (e.g., Somoza, 1998), and negative trench roll-back caused generalized extension across the region (see Kay and Copeland, this volume, chapter 9). As shown in Figure 7D, extension at this time was localized in the Coya Machalí intra-arc basin along the main Andes to the north (e.g., Godoy et al., 1999; Charrier et al., 2002), whereas extension near the Cortaderas lineament to the south was more intense, with effects extending from the Cura Mallín intra-arc basin (e.g., Burns et al., this volume, chapter 8) into the retroarc where alkali olivine basalts were erupting (see Ramos and Barbieri, 1989; Kay, 2001b; Kay et al., 2004; Kay and Copeland, this volume, chapter 9). The cartoon in Figure 7D shows a return to a steeper subduction zone than
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Tectonic evolution of the southern Central Andes, Mendoza and Neuquén
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that which existed in the Late Cretaceous to Paleocene in the southern region, which is in accord with extension in the Cura Mallín basin and into the backarc.
contractional deformation was less intense, the subduction zone for the southern segment in Figure 7E is shown as steeper (~30°) than that to the north.
Shallowing of Benioff Zone in Northern Sector (Middle to Late Miocene)
Steepening of the Benioff Zone (Pliocene to Quaternary)
As recognized by Kay (2001a, 2001b, 2002; Kay et al., this volume, chapter 2), an important transient shallowing of the subduction zone north of the Cortaderas lineament can explain the Miocene advance of deformation into the foreland and the eruption of volcanic rocks with arc-like magmatic signatures up to 500 km east of the trench. This shallowing, which is depicted in the cartoon in Figure 7E, fits well with the wave of deformation in the Malargüe fold-and-thrust belt and the foreland expansion of magmatism discussed already. The age and distribution of magmatic rocks in the Principal Cordillera and the synorogenic deposits in the foreland (Kraemer et al., 2000) are consistent with the eastward advance of the Malargüe thrust front from 15 Ma to 5 Ma as the magmatic arc broadened. During the time of shallow subduction, the San Rafael block and the Sierra de Chachahuén were uplifted in a manner analogous to that of the Sierras Pampeanas over the present Pampean (Chilean) flat-slab segment to the north. Because magmatism did not completely cease, the shallowing of the slab is considered to have been less pronounced than under the Pampean flat slab to the north (Kay, 2001b, 2002; Mancilla, 2001; Kay et al., this volume, chapters 2 and 10). During this same period, the region south of the Cortaderas lineament was subjected to contractional deformation that led to the middle to late Miocene compressional inversion of the Cura Mallín basin, reactivation of the Agrio fold-and-thrust belt, and the uplift of the Chihuidos high. As magmatism in the south was restricted to the region of the Principal Cordillera and
Figure 7. Conceptual cartoons (not to scale) comparing the tectonic evolution of the sectors north and south of the Cortaderas lineament: (A) Triassic–Early Jurassic rifting associated with the breakup of Pangea and inception of early Jurassic subduction. (B) Jurassic–Early Cretaceous subduction with generalized extension associated with negative trench roll-back. (C) Late Cretaceous to Paleogene contraction in the Agrio fold-and-thrust belt (FTB) and less extensive tectonic inversion in the Malargüe fold-and-thrust belt. (D) Extension associated with steepening of the subduction zone related to negative trench roll-back velocity. (E) Differential shallowing of the subduction zone with Miocene contractional deformation and crustal shortening. (F) Steepening of the subduction zone, widespread extension and magmatism in the northern sector; localized extension and magmatism, and initiation of contraction in the Guañacos fold-and-thrust belt. (G) Present setting under contraction. More detailed early Miocene to late Quaternary lithospheric scale cross sections north of the Cortaderas lineament are included in Kay et al. (this volume, chapter 2), and more detailed Pliocene to late Quaternary sections comparing the regions north and south of the Cortaderas lineament are in Folguera et al. (this volume, chapter 12).
The cartoons in Figures 7F and 7G show a return to a steeper subducting slab than in the late Miocene, with the most pronounced steepening in the north. Kay (2001b, 2002; Kay et al., this volume, chapter 2) argued that the end of arclike magmatism far east of the trench followed by widespread Pliocene to Quaternary mafic volcanism with a progressively more intraplate-like chemical signature is best explained by such a steepening of the subducting slab north of the Cortaderas lineament. The widespread within-plate basaltic volcanism of the Payenia volcanic field is thought to be triggered by the reinsertion of hot asthenosphere into the thicker mantle wedge above the steepening slab (Kay et al., 2004). Such a scenario also fits with the Pliocene to Quaternary extensional collapse of foreland in this paper. Crustal melting by basaltic underplating linked to the injection of hot asthenosphere along the cordilleran axis north of the Cortaderas lineament can explain the formation of large Quaternary collapse calderas and the emplacement of rhyolitic domes along the Las Loicas trough (Folguera et al., this volume, chapter 12). Crustal weakening would favor shortening along the PliocenePleistocene arc leading to the development of the Guañacos fold-and-thrust belt along the eastern slope of the Principal Cordillera (see Folguera et al., this volume, chapter 11). The region south of the Cortaderas lineament records a contraction of the Pliocene-Pleistocene arc, which retreated westward to the present position in the southern sector of the Southern Volcanic Zone (Lonquimay) volcanic arc. Pliocene to Quaternary extension in this region is confined to the Loncopué trough. Local negative trench roll-back has been postulated to explain both the extension and minor retreat of the volcanic arc (see Lara and Folguera, this volume, chapter 14). This roll-back could be accompanied by a minor steepening of the subducting slab (Folguera et al., this volume, chapter 12). Roll-back could also play a role in events in the northern sector, but only shallowing followed by steepening of the subducting slab can easily explain expansion followed by disappearance of arc-like magmatic activity and compressional deformation far into the backarc. CONCLUDING REMARKS The tectonic evolution of the southern Mendoza and Neuquén Andes is characterized by changes in the position of the magmatic arc front, periods of expansion of Tertiary to Holocene volcanism into the foreland, and waves of compression or extension that are best linked to changes in the geometry of the Benioff zone. In detail, shifting or expansion of arc magmatism into the foreland is coeval with contraction in the fore-
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land fold-and-thrust belt, whereas retraction of arc magmatism toward the trench is linked with extension and collapse. There is no clear evidence to suggest a correlation of changes in slab geometry with subduction of aseismic ridges, as has been proposed for the Chilean flat slab farther north (e.g., Yáñez et al., 2001; Ramos et al., 2002). The simplest explanation for these changes is that subduction of a ridge produced a transient shallowing in the geometry of the subducted slab (see also Kay et al., this volume, chapter 9). These changes could have been enhanced by the absolute motion of the South American plate as normally encompassed in the overriding velocity of the upper plate (Jarrard, 1986; Sobolev and Babeyko, 2005). The overriding velocity is equivalent to the trench roll-back velocity plus the orogenic shortening rate. An increase in the overriding velocity may result in compression and shortening and positive trench roll-back velocities as proposed by Daly (1989); a decrease in the overriding velocity may produce an extensional regime coeval with negative trench roll-back. Evaluation of the mechanisms that control changes in the tectonic regime in the Neuquén Andes requires taking into account changes that affect a broader area. For example, Miocene compression is well known all along the Andean margin. Silver et al. (1998) argued that the resulting contraction can be related to a relative increase in the overriding velocity of the entire South American plate as Africa slowed down. In another example, the extensional regime that controlled the inception of normal fault-bounded intra-arc basins in the Oligocene is well known to have occurred all along the Andean margin (e.g., Daly, 1989; Mpodozis and Ramos, 1989). This change could be attributed to a general decrease in the overriding velocity affecting the South America plate, which would produce a steepening of the subduction zone and a retreat of magmatic activity toward the trench. On the other hand, local conditions might have enhanced these effects, as seen in the southern sector of the Neuquén Andes, where widespread normal faults related to Mesozoic rifts played a major role in later contractional events. ACKNOWLEDGMENTS The authors acknowledge financial support from the Agencia Nacional de Promoción Científica y Tecnológica (grant ANCPYT 14144/03 to Ramos) and Repsol YPF (to S.M. Kay). Members of the Laboratorio de Tectónica Andina (Universidad de Buenos Aires) and the Cornell Andes Group are thanked for interesting discussions and comments. The paper was improved by careful reviews by Constantino Mpodozis and Robert Kay. REFERENCES CITED Bastías, H., Tello, G.E., Perucca, L.P., and Paredes, J.D., 1993, Peligro sísmico y neotectónica, in Ramos, V.A., ed., Geología y recursos naturales de Mendoza: Buenos Aires, XII Congreso Geológico Argentino y II Congreso de Exploración de Hidrocarburos, Relatorio VI-1, p. 645–658. Bermúdez, A., 1991, Sierra del Nevado. El límite oriental del arco volcánico Neógeno entre los 35°30′ y 36° L.S. Argentina: 6th Congreso Geológico Chileno Actas, v. 1, p. 318–322.
Bermúdez, A., Delpino, D., Frey, F., and Saal, A., 1993, Los basaltos de retroarco extraandinos, in Ramos, V.A., ed., Geología y recursos naturales de Mendoza: Buenos Aires, XII Congreso Geológico Argentino y II Congreso de Exploración de Hidrocarburos, Relatorio I-13, p. 161–172. Burckhardt, C., 1900, Profils géologiques transversaux de la Cordillère Argentino-Chilienne: Museo de La Plata, Anales, Sección Geología y Mineralógica, v. 2, p. 1–136. Burns, W.M., 2002, Tectonic and depositional evolution of the Tertiary Cura Mallín Basin in the southern Andes (36.5 to 38°S lat.) [Ph.D. Thesis]: Cornell University, Ithaca, New York, 218 p. Burns, W.M., Jordan, T.E., Copeland, P., and Kelley, S.A., this volume, The case for extensional tectonics in the Oligocene-Miocene Southern Andes as recorded in the 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, doi: 10.1130/2006.2407(08). Charrier, R., Baeza, O., Elgueta, S., Flynn, J.J., Gans, P., Kay, S.M., Muñoz, 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, no. 1, p. 117–140. Chernicoff, J., and Zappettini, E., 2004, Geophysical evidence for terrane boundaries in south-central Argentina: Gondwana Research, v. 7, p. 1105–1117, doi: 10.1016/S1342-937X(05)71087-X. Cisternas, M.E., and Frutos, J., 1994, Evolución tectono-estratigráfica de la Cuenca Terciaria de los Andes del Sur de Chile (37°30′–40°30′Lat.S.): 7th Congreso Geológico Chileno (Concepción) Actas, v. 1, p. 6–12. Cobbold, P.R., and Rossello, E.A., 2003, Aptian to Recent compressional deformation of the Neuquén Basin, Argentina: Marine and Petroleum Geology, v. 20, p. 429–443, doi: 10.1016/S0264-8172(03)00077-1. Cobbold, P.R., Diraison, M., and Rossello, E.A., 1999, Bitumen veins and Eocene transpression, Neuquén Basin, Argentina: Tectonophysics, v. 314, p. 423–442, doi: 10.1016/S0040-1951(99)00222-X. Corbella, H., Novas, F.E., Apesteguía, S., and Leanza, H.A., 2004, First fission track-age for the dinosaur-bearing Neuquén Group (Upper Cretaceous) Neuquén Basin, Argentina: Revista Museo Argentino de Ciencias Naturales (N.S.), v. 6, p. 1–6. Costa, C.H., Cisneros, H., Salvarredi, J., and Gallucci, A., 2004, Nuevos datos y reconsideraciones sobre la neotectónica del margen oriental del bloque de San Rafael: 12 Reunión Sobre Microtectónica y Geología Estructural (Cafayate), Resúmenes, p. 7. Dajczgewand, D.M., 2002, Faja plegada y corrida de Malargüe: Estilo de deformación en la región de Mallín Largo. Trabajo Final de Licenciatura: Buenos Aires, Universidad de Buenos Aires (unpublished), 119 p. Daly, M., 1989, Correlations between Nazca/Farallón plate kinematics and forearc evolution in Ecuador: Tectonics, v. 8, p. 769–790. Delpino, D.H., and Bermúdez, A., 1985, Volcán Plateado. Vulcanismo andesítico de retroarco en el sector extrandino de la Provincia de Mendoza, 35°42′ Lat. Sur. Argentina: Antofagasta, 4th Congreso Geológico Chileno Actas, v. 3, p. 108–119. Dungan, M.A., Wulff, A., and Thompson, R., 2001, Eruptive stratigraphy of the Tatara–San Pedro complex, 36°S, Southern Volcanic Zone, Chilean Andes: Reconstruction method and implications for magma evolution at long-lived arc volcanic centers: Journal of Petrology, v. 42, p. 555–626, doi: 10.1093/petrology/42.3.555. Folguera, A., and Ramos, V.A., 2000, Control estructural del Volcán Copahue: Implicancias tectónicas para el arco volcánico Cuaternario (36°–39°S): Asociación Geológica Argentina Revista, v. 55, p. 229–244. Folguera, A., and Ramos, V.A., 2001, Distribución de la deformación en los Andes Australes (33°–46°S), in Cortés, J.M., Rossello, E., and Dalla Salda, L., eds., Avances en microtectónica: Asociación Geológica Argentina, Serie D, Publicación Especial, v. 5, p. 13–18. Folguera, A., and Ramos, V.A., 2002, Partición de la deformación durante el Neógeno en los Andes Patagónicos Septentrionales (37°–46°S): Revista de la Sociedad Geológica de España, v. 15, p. 81–93.
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Tectonic evolution of the southern Central Andes, Mendoza and Neuquén Folguera, A., Ramos, V.A., and Melnick, D., 2002, Partición de la deformación en la zona del arco volcánico de la cordillera Neuquina en los últimos 30 millones de años (36°–39°S): Revista Geológica de Chile, v. 29, p. 227–240. Folguera, A., Ramos, V.A., and Melnick, D., 2003, Recurrencia en el desarrollo de cuencas de intraarco. Colapso de estructuras orogénicas. Cordillera Neuquina (37°30′): Revista de la Asociación Geológica Argentina, v. 58, p. 3–19. Folguera, A., Ramos, V.A., Hermanns, R.L., and Naranjo, J., 2004, Neotectonics in the foothills of the southernmost central Andes (37°–38°S): Evidence of strike-slip displacement along the Antiñir-Copahue fault zone: Tectonics, v. 23, TC5008, doi: 10.1029/2003TC001533. Folguera, A., Folguera, A., Zárate, M., and Ramos, V.A., 2005, La cuenca de antepaís Neógena del Río Negro asociada con el levantamiento de los Andes de Neuquén, in 16th Congreso Geológico Argentino Actas, La Plata: Actas, v. 2, p. 29–36. Folguera, A., Ramos, V.A., González Díaz, E.F., and Hermanns, R., 2006, this volume, Miocene to Quaternary deformation of the Guañacos fold-andthrust 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°S lat): Geological Society of America Special Paper 407, doi: 10.1130/ 2006.2407(11). Folguera, A., Zapata, T., and Ramos, V.A., 2006, this volume, Late Cenozoic extension and the evolution of the Neuquén 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, doi: 10.1130/ 2006.2407(12). Franchini, M.B., López Escobar, L., Shalamuk, I.B.A., and Meinert, L.D., 2003, Paleocene calc-alkaline subvolcanic rocks from Nevazón Hill area (NW Chos Malal fold belt), Neuquén, Argentina, and comparison with granitoids of the Neuquén-Mendoza volcanic province: Journal of South American Earth Sciences, v. 16, p. 399–422, doi: 10.1016/S08959811(03)00103-2. Franzese, J.R., and Spalletti, L.A., 2001, Late Triassic continental extension in southwestern Gondwana: Tectonic segmentation and pre-break-up rifting: Journal of South American Earth Sciences, v. 14, p. 257–270, doi: 10.1016/S0895-9811(01)00029-3. García Morabito, E., 2005, Geología del sector occidental de la depresión de Loncopúe, provincia del Neuquén: Buenos Aires, Trabajo Final de Licenciatura, Universidad de Buenos Aires (inédito), 108 p. Gerth, E., 1931, La estructura geológica de la Cordillera Argentina entre el Río Grande y el Río Diamante en el sud de la provincia de Mendoza: Academia Nacional de Ciencias Actas, v. X, p. 125–172. 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. 217–236, doi: 10.1016/S0040-1951(99) 00046-3. González Díaz, E.F., 1964, Rasgos geológicos y evolución geomorfológica de la Hoja 27d (San Rafael) y zona occidental vecina (Provincia de Mendoza): Asociación Geológica Argentina, Revista, v. 19, p. 151–188. González Díaz, E.F., 1972a, Descripción geológica de la Hoja 27d San Rafael, Provincia de Mendoza: Servicio Nacional Minero Geológico, Boletín, v. 132, p. 1–127. González Díaz, E.F., 1972b, Descripción geológica de la Hoja 30d PayúnMatrú, Provincia de Mendoza: Dirección Nacional de Geología y Minería, Boletín, v. 130, p. 1–88. Gräfe, K., Glodny, J., Seifert, W., Rosenau, M., and Echtler, H., 2002, Apatite fission track thermochronology of granitoids at the south Chilean active continental margin (37º–42ºS): Implications for denudation, tectonics and mass transfer since the Cretaceous, in Proceedings of the 5th International Symposium on Andean Geodynamics, Toulouse, France: IRD Editions, Paris, , p. 275–278.
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MANUSCRIPT ACCEPTED BY THE SOCIETY 22 DECEMBER 2005
Printed in the USA
Geological Society of America Special Paper 407 2006
Upper Cretaceous to Holocene magmatism and evidence for transient Miocene shallowing of the Andean subduction zone under the northern Neuquén Basin Suzanne Mahlburg Kay* W. Matthew Burns*† Institute for the Study of the Continents and Department of Earth and Atmospheric Sciences, Snee Hall, Cornell University, Ithaca, New York 14853, USA Peter Copeland* Department of Geosciences, University of Houston, Houston, Texas 77204, USA Oscar Mancilla* Repsol YPF, Buenos Aires, Argentina
ABSTRACT Evidence for a Miocene period of transient shallow subduction under the Neuquén Basin in the Andean backarc, and an intermittent Upper Cretaceous to Holocene frontal arc with a relatively stable magma source and arc-to-trench geometry comes from new 40Ar/ 39Ar, major- and trace-element, and Sr, Pb, and Nd isotopic data on magmatic rocks from a transect at ~36°–38°S. Older frontal arc magmas include early Paleogene volcanic rocks erupted after a strong Upper Cretaceous contractional deformation and mid-Eocene lavas erupted from arc centers displaced slightly to the east. Following a gap of some 15 m.y., ca. 26–20 Ma mafic to acidic arclike magmas erupted in the extensional Cura Mallín intra-arc basin, and alkali olivine basalts with intraplate signatures erupted across the backarc. A major change followed as ca. 20–15 Ma basaltic andesite–dacitic magmas with weak arc signatures and 11.7 Ma Cerro Negro andesites with stronger arc signatures erupted in the near to middle backarc. They were followed by ca. 7.2–4.8 Ma high-K basaltic to dacitic hornblendebearing magmas with arc-like high field strength element depletion that erupted in the Sierra de Chachahuén, some 500 km east of the trench. The chemistry of these Miocene rocks along with the regional deformational pattern support a transient period of shallow subduction that began at ca. 20 Ma and climaxed near 5 Ma. The subsequent widespread eruption of Pliocene to Pleistocene alkaline magmas with an intraplate chemistry in the Payenia large igneous province signaled a thickening mantle wedge above a steepening subduction zone. A pattern of decreasingly arc-like Pliocene
*E-mails: Kay—
[email protected]; Burns—
[email protected]; Copeland—
[email protected]; Mancilla—
[email protected]. †Now at U.S. Geological Survey, Reston, Virginia 20192, USA.
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, doi: 10.1130/2006.2407(02). For permission to copy, contact
[email protected]. ©2006 Geological Society of America. All rights reserved.
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S.M. Kay et al. to Holocene backarc lavas in the Tromen region culminated with the eruption of a 0.175 ± 0.025 Ma mafic andesite. The northwest-trending Cortaderas lineament, which generally marks the southern limit of Neogene backarc magmatism, is considered to mark the southern boundary of the transient shallow subduction zone. Keywords: Andes, volcanism, tectonics, Neuquén Basin, shallow subduction, geochemistry, Neogene.
INTRODUCTION The history of the continental lithosphere, mantle wedge, and subducting plate of the south Central Andes, between 36.5°S and 38°S, is reflected in the temporal and spatial distribution and chemistry of the Upper Cretaceous to Holocene arc and backarc magmatic rocks of the Neuquén Basin. Together with the structural and geophysical characteristics of the region, the distribution and geochemical features of these magmatic rocks can be used to formulate a model for the magmatic and deformational history and the evolution of the subducting slab in a west to east transect between 36° and 37.5°S through the Neuquén Basin. GEOLOGIC AND TECTONIC SETTING The east-west transect through the Neuquén Basin between 36.5° and 38°S latitude lies just south of where the Holocene volcanic arc front of the Southern Volcanic Zone is displaced to the west (Fig. 1). At this latitude, the currently subducting Nazca plate corresponds to chrons 8–13 (ca. 33–25 Ma; Cande and Kent, 1992). The Holocene centers of the transitional Southern Volcanic Zone segment to the north dominantly erupt andesitic lavas, whereas those of the southern Southern Volcanic Zone segment dominantly erupt high-Al basalts (see review by Stern, 2004). To the east, the backarc can be divided into two regions by the northwest-trending Cortaderas lineament (Fig. 2), which broadly intersects the southern end of the transitional Southern Volcanic Zone segment. North of the Cortaderas lineament, Miocene to Holocene backarc magmatic rocks are widespread in a retroarc region where Mesozoic rifting was less important and Neogene contractional deformation was more important than to the south (see Ramos and Kay, this volume, chapter 1). Particularly notable in the backarc are the Pleistocene to Holocene backarc Tromen and Payún Matrú volcanic centers and the extensive mafic flows that constitute the Payenia and Auca Mahuída volcanic fields (Figs. 1 and 2). To the south of the Cortaderas lineament, Miocene to Holocene backarc magmatic rocks are essentially absent. The ages and locations of the Upper Cretaceous to Holocene magmatic rocks discussed in this paper are summarized in Table 1. They largely overlie or intrude the Mesozoic to early Paleogene sedimentary strata of the Neuquén Basin. The history of the Neuquén Basin can be divided into three general stages (e.g., Vergani et al., 1995): (1) a Triassic to Early Jurassic
Figure 1. Generalized map of Eocene to Holocene magmatic rocks of the Andean Cordillera and Patagonia from 34° to 52°S, modified from the 1:2,500,000 scale, 1997 geologic map of Argentina (Servicio Geológico Minero Argentino, Buenos Aires) and map in Stern et al. (1990). Boxed area labeled “Central Neuquén Basin transect” between ~36.5 and 38°S is region considered in this paper. SVZ—Southern Volcanic Zone.
Upper Cretaceous to Holocene magmatism
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Figure 2 (on this and following page). Maps of (A) pre-Pliocene and (B) post-Miocene magmatic units in Argentina erupted over the Neuquén Basin region. Maps are modified from the provincial maps of Neuquén (Delpino and Deza, 1995) and Mendoza (Servicio Geológico Minero Argentino, Buenos Aires). Cortaderas lineament is after Ramos (1978). Circles in early Miocene units in backarc schematically indicate late early Miocene volcanic centers. Labels in the undifferentiated Miocene volcanic fields are for the early Miocene Cura Mallín Formation (cm) and the Miocene Trapa Trapa (tt), Cajón Negro (cn), Quebrada Honda (qh) Formations, and the late Miocene to Pliocene? Pichi Neuquén Complex (pn). PQ—Plio-Quaternary. Holocene volcanic centers of the Southern Volcanic Zone arc are from geologic maps of Chile and Argentina.
prerift and rift stage, (2) an Upper Jurassic to Cretaceous subsidence stage, and (3) a Paleocene to Holocene modification stage. The first stage was largely shaped by the extension and fault-controlled subsidence that preceded and accompanied the initial breakup of the Pangea supercontinent. The widespread Triassic Choiyoi rhyolitic volcanic rocks (e.g., Kay et al., 1989) exposed in the Cordillera del Viento (Fig. 2) and underlying much of the Neuquén Basin erupted at this time. The active rifting of this stage generally terminated as Middle Jurassic Andean tectonism and magmatism began to the west. During the second stage, the discrete rifts and intervening basement blocks of the first stage generally merged into a broad postrift basin that was filled by Middle Jurassic to Paleogene sedimentary strata. In the last stage, the Neuquén Basin strata were modified by Tertiary to Holocene extensional and contractional
deformation and affected by periodic magmatic events across the basin. The magmatic rocks of this third stage are the principal topic of this paper. Distribution, Age, and Chemistry of Neuquén Basin Magmatic Rocks The distribution and ages of the Upper Cretaceous to Holocene Neuquén Basin magmatic rocks described below are shown on maps in Figures 1–4 and summarized in Table 1. Twelve new 40Ar/ 39Ar ages are listed in Table 2. Age spectra are presented in Appendix 1. Analytical techniques are the same as in Jordan et al. (2001). New major- and trace-element data for 90 samples and isotopic data for 12 samples are listed in Tables 3–7 and plotted along with 300 other unpublished and pub-
22
S.M. Kay et al.
Figure 2 (continued).
TABLE 1. ARC AND BACKARC MAGMATIC UNITS FROM THE NEUQUÉN ANDES AND NEUQUÉN BASIN (36.5° TO 38°S LATITUDE) Age Arc and near arc Near to middle backarc Far backarc (Ma) Quaternary/Holocene <0.5 Southern Volcanic Zone Tromen Center Cerro Tromen andesite Late Pliocene/Quaternary 2–1 Southern Volcanic Zone Chos Malal trough, etc. Auca Mahuida, Payún Matrú fields Guañacos Formation basalts (Chapúa, Maipo, El Puente Fms.) 2.3 (Cerro Tanque) Pliocene 3.6–2.6 Centinela Formation 4–4.5 Cola de Zorro, Pichi Neuquén Formation Tromen Tilhué Formation rhyolites basalts west of Payún Matrú field Tromen Coyocho Formation flows 4.5 Parva Negra basalt late Miocene 7.5–4.8 Cajon Negro Formation? Huincán II Formation (Mendoza) Chachahuén volcanic complex Cajon Negro Formation? 11.7 Cerro Negro andesite middle Miocene 14–11 Trapa Trapa Formation Huincán I Formation (Mendoza) Pichi Tril Formation early middle Miocene 20–15 Trapa Trapa Formation and plutons Sierras Huantraico/Negra centers Molle Formation (Mendoza) early Miocene 25–20 Cura Mallín Formation Sierra de Huantraico alkali basalts Matancilla region alkali basalts Eocene 50 ± 5 Cerro Bayo de la Esperanza 50–45 Caicayén group Paleogene/early Eocene 56–50 Cayanta Formation 60–56 Cerro Nevazón region Latest Cretaceous 75–66 Varvarcó Pluton Pelán Formation Period
Upper Cretaceous to Holocene magmatism
23
Figure 3. Map of Tertiary to Holocene magmatic rocks in the Tromen massif and Chos Malal trough region modified from the northwestern portion of the 1:200,000 scale Buta Ranquil geological map (Hoja 32c) of Zollner and Amos (1973) and the northeastern portion of the 1:200,000 scale Chos Malal geological map (Hoja 32b) of Holmberg (1976). Pliocene to Holocene mapping units are those of Holmberg (1976). Ages on map are 40Ar/ 39Ar ages in Table 2 and Kay and Copeland (this volume, chapter 9). White dots are locations of samples listed in Table 6 and plotted in Figures 5–10.
lished analyses in the fields in Figures 5–9. Sources of published data are in the figure captions. Analytical techniques are as described in Kay et al. (this volume, chapter 10). Sample locations are listed in Appendix 2. The range of magmatic rock types in the Neuquén Basin region is illustrated in the SiO2 versus total alkali (Na2O + K2O) concentration diagrams in Figure 5. Major-element, trace-element, and isotopic data are used in the following to characterize the Neuquén Basin magmatic rocks. Characteristics of basaltic to mafic andesitic samples, particularly those with low FeO/MgO ratios (~0.7–1.0) and high Cr and Ni (>200–300 ppm) concentrations, are useful in interpreting mantle and subcrustal processes, whereas characteristics of silicic andesite to rhyolites are useful in interpreting crustal processes. Ratios and concentrations of incompatible elements (concentrated in melts) provide insights into mantle magma sources and tectonic settings. Slab-related processes are reflected in ratios of Ti group elements (Ta, Nb) to rare earth elements (REEs) and alkali (K, Rb, Cs)–alkaline earth (Ba, Sr) elements. Ratios of La/Ta, Ba/La, and Ba/Ta in Neuquén Basin samples are compared with those of Southern Volcanic Zone frontal arc magmas (La/Ta ~ 40–95; Ba/La > 20; and Ba/Ta >
500) and mid-ocean-ridge basalt (MORB) and oceanic-island basalt (OIB) magmas (La/Ta < 12; Ba/La < 15) from Hickey et al. (1986) in Figure 6. Indicators of mantle source conditions include the high field strength elements (HFSEs) plotted in Figure 7. High Ta/Hf ratios reflect an enriched intraplate mantle source, whereas low Ta/Hf ratios indicate depleted MORB or arc mantle sources. High Th/Hf ratios are typical of calc-alkaline arc sources. Relative concentrations of incompatible elements also serve as guides to percentages of partial melting in mantle and crustal source regions. Ratios and concentrations of compatible elements (concentrated in minerals) reflect residual minerals that are either left in the magma source after melting or removed by fractionation processes. The residual mineral assemblage reflects the pressure, temperature, and fluid conditions under which the magma last equilibrated. Trace elements are useful in determining residual mineral assemblages because: (1) olivine, orthopyroxene, and micas have little affinity for REEs, but take Ni and Cr, (2) feldspar takes Eu2+ and Sr, (3) clinopyroxene, and to a greater extent amphibole, take middle and heavy REEs and Sc, (4) garnet takes heavy REEs, and (5) accessory titanite and apatite take middle
24
S.M. Kay et al.
Figure 4. Generalized Holocene magmatic and sedimentary rocks west of the Cordillera del Viento and east of the Southern Volcanic Zone (SVZ). Names and locations (points) are shown for samples with analyses in Tables 3–5. Map is after Burns (2002).
Upper Cretaceous to Holocene magmatism
25
TABLE 2. NEW 40Ar/39Ar GEOCHRONOLOGY FOR NEUQUÉN BASIN MAGMATIC ROCKS (SPECTRA IN APPENDIX 1) 40 Ar/39Ar age Sample Unit and locality Type (±1 sigma) (Ma) BPN11 Varvarcó pluton - hornblende-bearing leucogranodiorite Biotite 69.09 ± 0.13 West side of road, east of river, north of Varvarco (36°49.42 S, 70°40.41 W) TDR21 Cerro Negro, andesite flow Hornblende 11.70 ± 0.20 Puesto on southeast side of Cerro Negro (36°50.55 S, 69°57 W) Chos Malal trough and Tromen Massif Biotite 4.0 ± 0.4 TDR16 Tilhué Formation, Cerro Bayo de Tromen - rhyolite dome (37°41.8 S, 69°34 W) TDR19 Intermediate Group Chapúa basalt with large plagioclase phenocrysts Groundmass 1.44 ± 0.08 Road cut north of Chapúa School ( 37°10.77 S, 70°14.95 W) TDR6 Cerro Waile andesite, El Puente Fm. Groundmass 1.04 ± 0.06 1/3 of way up on road to top of Cerro Waile (37°03.6 S, 70°09.5 W) TDR2 Tromen "escorial" andesitic flow Groundmass 0.175 ± 0.028 Cerro Tromen north of Laguna Tromen (37°05.50 S, 70°05.5 W) Southern Payun Matru Field DR38 Flow from Cerro Tanque, Sierra de Chachahuén Northwest of Cerro Campanario (37°0.54 S, 68°56.64' W) DRC14 Young basal flow from well preserved Cerro Méndez cone Just north of Rio Colorado (37°19.45 S, 68°57.30 W) Auca Mahuida Field RD3 Plateau basalt flow, Auca Mahuida West of Puesto Agua del Macho (37°56'S, 69°06'W) Pampa de Las Yeguas basalt in Ardolino et al. (1996) RD1 Plateau basalt flow, Auca Mahuida East of Total well in Rincón Chico (37°49.5 S, 69°46 W), Cerro Las Liebres basalt in Ardolino et al. (1996) RD8 Plateau basalt flow, eastern Auca Mahuida West of the El Cruce Gomeria (37°41.79 S, 68°27.82 W) Cerro Grande basalt in Ardolino et al. (1996) RD20b Mugearite/benmorite flow from crater rim, Auca Mahuida Quebrada on southeast side of rim (37°46.04 S, 68°54.15 W) Auca Mahuida Group of Ardolino et al. (1996)
REEs and zircon takes heavy REEs, Hf, Th, and U. Increasing pressure can produce a change from pyroxene to amphibole to garnet in the residual mineralogy that can be detected by increasing La/Yb and Sm/Yb ratios. La/Yb ratios provide a guide to the overall steepness of the REE pattern (Fig. 8A), whereas La/Sm and Sm/Yb ratios (Fig. 8B) provide a guide to light and heavy REE behavior. Nd, Sr, and Pb isotopic ratios (Figs. 9 and 10) contain independent source region information because they reflect parent/daughter ratios in closed systems and contaminant addition in open systems. Magmas with higher 87Sr/ 86Sr, 206Pb/ 204Pb, 207Pb/ 204Pb, and 208Pb/ 204Pb and lower 143Nd/ 144Nd ratios are said to be relatively isotopically “enriched.” Upper Cretaceous to Eocene Magmatic Rocks The Upper Cretaceous to Paleogene magmatic history of the Neuquén Basin is not well known. Volcanic and plutonic rocks of this age generally occur along a north-south–trending belt that runs through northwestern Neuquén (Fig. 2A). Unlike younger magmatic rocks, they occur only in the western Neuquén Basin. Radiometric ages support dividing them into: (1) Upper Cretaceous, (2) Paleocene, and (3) latest Paleocene to Eocene groups. All intrude or overlie deformed Mesozoic strata, but are not themselves significantly deformed (Llambías et al., 1978; Llambías and Rapela, 1989; Franchini et al., 2003).
Groundmass
2.07 ± 0.11
Groundmass
1.23 ± 0.17
Groundmass
1.78 ± 0.1
Groundmass
1.55 ± 0.07
Groundmass
1.39 ± 0.14
Groundmass
0.99 ± 0.04
Upper Cretaceous to Paleocene magmatism has been summarized by Franchini et al. (2003). Upper Cretaceous activity is confirmed by K/Ar ages of 74.2 ± 1.4 Ma for a biotite from an andesite dike in the Campana Mahuída region (Sillitoe, 1977), a whole-rock K/Ar age of 71.5 ± 5 Ma for an amphibole-bearing andesite sill cutting the Pelán Unit in the Cerro Nevazón region on the east side of the Cordillera del Viento (Llambías et al., 1978; Linares and González, 1990), whole-rock K/Ar ages from 69 ± 4 to 65 ± 3 Ma for tuffs and veins between Andacollo and Huinganco (Vilas and Valencio, 1978), and a whole-rock age of 67 ± 3.2 Ma for a tonalite stock near El Maitenes in the southern Cordillera del Viento (Domínguez et al., 1984). Franchini et al. (2003) also referred to an unpublished age of 64.7 ± 3.2 Ma for a pluton in the Cordillera del Viento. In addition, Zamora Valcarce et al. (this volume, chapter 6) report 40Ar/ 39Ar ages of 72.83 ± 0.83 Ma and 65.5 ± 0.46 Ma, respectively, for a volcanic bomb and the Cerro Naunauco laccolith in the Collipilli region. A new 40Ar/ 39Ar biotite age of 69.09 ± 0.13 Ma from a granodiorite pluton near Varvarcó in the western Cordillera del Viento in Table 2 is interpreted as a cooling age. Paleocene magmatism is confirmed by a 40Ar/ 39Ar age of 56.64 ± 0.44 Ma for an andesitic sill in the Collipilli region (Zamora Valcarce et al., this volume, chapter 6) and by hornblende K/Ar ages of 59.1 ± 2.9 Ma and 56.5 ± 1.7 Ma from a gabbro and a diorite in the Cerro
5.05 2.10 4.5 1.5 6.6 0.82 31.6 876 27.8 0.34 0.147 1.38 135 4961 3.5 36°49.42 70°40.41 1265 Cret.
13.9 32.1 13.1 3.12 0.796 0.448 2.10 0.295 361 439 3.3 1.3 4.7 3.4 0.50 12.1 5 2 10
La Ce Nd Sm Eu Tb Yb Lu Sr Ba Cs U Th Hf Ta Sc Cr Ni Co
Na2O+K2O FeO/MgO La/Sm Sm/Yb La/Yb Eu/Eu* Ba/La Ba/Ta La/Ta Th/La Ta/Hf Th/Hf Ba/Cs K/Cs Th/U Latitude °S Longitude °W Elevation (m) Age (Ma)
BPN11 67.24 0.46 14.34 4.74 0.13 2.26 5.67 3.11 1.95 0.06 99.95
Fm./Group Locality Sample SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O P2O5 Total
1 Varvarcó
4.50 2.05 3.9 1.7 6.7 0.82 22.5 1205 53.5 0.32 0.121 2.06 437 8687 4.1 37.139° 70.699°
15.8 33.0 16.2 4.02 1.17 0.626 2.37 0.332 513 356 0.8 1.2 5.0 2.4 0.30 20.9 9 6 18
Paleocene/Early Eocene
3.90 1.73 3.3 1.6 5.2 0.82 27.2 1091 40.1 0.25 0.119 1.19 489 7354 3.3 37.188° 70.691°
11.3 24.4 13.1 3.43 1.04 0.621 2.19 0.316 437 308 0.6 0.9 2.8 2.4 0.28 21.3 4 4 19 4.35 2.17 2.9 1.5 4.4 0.82 26.1 1211 46.3 0.25 0.096 1.10 662 17913 3.7 37.142° 70.729°
11.1 22.4 15.7 3.87 1.06 0.686 2.55 0.370 365 291 0.4 0.7 2.8 2.5 0.24 30.9 5 4 47 4.57 1.78 3.8 1.6 6.2 0.82 26.2 1122 42.8 0.29 0.110 1.37 20 296 3.5 37.145° 70.667°
14.4 30.7 15.3 3.80 1.05 0.626 2.33 0.330 375 376 19.1 1.2 4.2 3.1 0.34 36.6 21 1 23 4.51 2.04 3.6 1.6 5.7 0.82 20.9 981 47.0 0.28 0.103 1.36 257 7251 3.9 37.135° 70.711°
13.7 30.0 12.9 3.83 1.15 0.633 2.43 0.349 424 287 1.1 1.0 3.9 2.8 0.29 22.4 3 0 20 6.36 2.02 4.1 1.7 7.0 0.82 22.2 952 42.8 0.25 0.114 1.21 385 7362 4.1 37.118° 70.691°
30.5 67.4 35.8 7.41 2.03 1.031 4.38 0.614 673 677 1.8 1.8 7.5 6.2 0.71 13.5 4 4 12 6.08 2.36 3.1 1.9 5.9 0.82 20.1 1275 63.5 0.29 0.070 1.29 1307 20934 3.7 37.279° 70.736°
40.7 99.7 51.0 13.31 3.08 2.012 6.94 1.058 617 819 0.6 3.2 11.8 9.2 0.64 54.6 6 9 29
Eocene
7.95 2.65 3.7 1.9 6.9 0.82 20.7 1046 50.4 0.19 0.086 0.83 280 9561 3.8 37.270° 70.818°
21.7 52.5 26.1 5.84 1.65 0.836 3.15 0.461 525 450 1.6 1.1 4.1 5.0 0.43 13.1 2 5 12 6.41 1.82 2.5 2.9 7.4 0.97 27.6 1902 68.9 0.16 0.075 0.84 26 619 4.5 36°42 69°50.2
13.9 34.6 22.7 5.47 1.55 0.730 1.89 0.269 604 383 14.7 0.5 2.2 2.7 0.20 29.8 94 40 31 7.67 2.46 3.5 3.1 10.9 0.97 25.6 1949 76.2 0.19 0.091 1.33 48 1475 3.7 36°41 69°50.45
21.5 49.4 28.0 6.17 1.55 0.760 1.97 0.283 653 549 11.5 1.1 4.1 3.1 0.28 23.3 10 8 25
7.48 2.32 3.5 3.0 10.6 0.97 35.9 2589 72.2 0.20 0.094 1.33 74 2030 3.5 36°42 69°50.2
19.9 43.8 25.3 5.59 1.56 0.729 1.87 0.305 780 712 9.6 1.1 3.9 2.9 0.28 22.5 13 6 26
8.18 2.74 6.7 4.0 26.6 0.97 33.4 2313 69.3 0.41 0.157 4.49 369 4967 4.4 36°41 69°50.45
55.0 108.7 45.4 8.25 1.88 0.363 2.07 0.285 1402 1835 5.0 5.2 22.8 5.1 0.79 4.8 3 2 7
8.14 2.98 6.1 3.1 18.9 0.97 37.9 2437 64.3 0.45 0.157 4.54 1400 15159 4.2 36°50.55 69°57
43.0 86.8 37.6 7.01 1.53 0.571 2.28 0.348 959 1631 1.2 4.6 19.3 4.3 0.67 10.2 26 8 12
TABLE 3. WHOLE-ROCK CHEMISTRY OF LATE CRETACEOUS TO MIOCENE MAGMATIC ARC ROCKS FROM THE NEUQUÉN ANDES 2 3 4 5 6 7 8 9 10 11 12 13 14 Eocene Cayanta Formation west of the Cordillera del Viento late Paleocene to Eocene in southern Mendoza Between Andacolla and Cayanta Río Guañacos Cerro Bayo de la Esperanza ESA-2 ESA-7 ESA-12 ESA-11 ESA-5 ESA-3 RG-10 RG-7 TDR8a TDR9a TDR9c TDR9b TDR11 50.36 52.98 52.81 53.58 54.21 57.42 55.62 55.80 48.04 49.11 50.35 57.02 54.33 0.66 0.88 1.14 0.88 0.85 0.67 1.39 0.69 1.25 1.22 1.08 0.61 0.73 18.98 18.90 18.27 19.05 17.30 17.95 15.98 19.40 18.06 19.65 19.08 18.12 17.72 7.94 7.83 8.23 7.64 8.76 6.07 8.87 5.80 10.74 9.68 9.06 5.43 6.37 0.21 0.21 0.21 0.17 0.22 0.21 0.29 0.20 0.16 0.13 0.14 0.20 0.22 4.58 3.81 3.79 4.28 4.29 3.00 3.76 2.19 5.89 3.93 3.91 1.98 2.13 8.94 10.58 8.76 9.49 9.56 7.53 7.38 7.69 7.92 7.72 7.41 7.71 10.05 3.34 3.65 3.40 3.89 3.54 4.80 4.50 6.10 5.31 5.62 5.12 5.20 6.01 0.56 0.85 0.95 0.68 0.97 1.56 1.58 1.85 1.10 2.05 2.36 2.98 2.13 0.05 0.22 0.12 0.13 0.21 0.33 0.39 0.28 0.43 0.30 0.29 0.31 0.31 95.61 99.91 97.68 99.80 99.91 99.54 99.77 100.00 98.90 99.41 98.79 99.56 100.01
(continued)
5.80 1.31 4.5 4.7 20.8 0.97 16.0 282 17.6 0.19 0.441 1.47 543 13291 5.0 36°50 69°03
29.1 61.4 30.9 6.51 1.754 0.848 1.40 0.179 955 466 0.9 1.1 5.5 3.7 1.65 18.9 203 74 34
Las Lajas DRC16 49.93 1.96 16.75 8.49 0.16 6.50 8.73 4.43 1.37 0.76 99.08
15
16.0 36.2 21.6 3.72 0.873 0.470 1.70 0.239 577 529 1.8 1.3 3.7 2.5 0.36 6.1 2 2 5
7.32 17.8 11.8 2.82 0.896 0.473 1.78 0.285 651 451 5.6 1.2 3.2 2.1 0.26 10.2 5 3 3
14.2 30.3 12.0 3.16 1.070 0.460 1.77 0.252 647 1436 11.4 1.4 3.3 2.3 0.29 11.4 7 2 8
10.3 23.8 12.4 2.76 0.733 0.362 1.62 0.226 425 957 8.7 1.0 4.1 2.4 0.28 9.0 4 2 5
19.6 43.9 29.3 6.23 1.38 1.06 2.88 0.416 422 477 0.6 1.4 5.1 4.0 0.52 28.6 38 18 26
11.6 25.5 12.7 3.55 1.04 0.583 1.89 0.279 384 235 0.5 0.8 2.9 2.3 0.25 28.5 43 24 30
37.8 87.7 44.7 11.7 2.06 2.24 5.65 0.881 368 633 5.4 3.7 13.4 8.4 0.97 34.8 18 15 28
5.31 8.33 4.70 4.84 Na2O+K2O FeO/MgO 2.64 1.33 1.82 3.17 La/Sm 4.3 2.6 4.5 3.7 3.1 3.3 3.2 Sm/Yb 2.2 1.6 1.8 1.7 2.2 1.9 2.1 La/Yb 9.4 4.1 8.1 6.4 6.8 6.1 6.7 Eu/Eu* 0.79 0.97 1.09 0.51 0.68 0.91 0.51 Ba/La 33.0 61.6 100.9 92.8 24.4 20.3 16.8 Ba/Ta 1486 1755 4989 3439 912 953 649 La/Ta 45.0 28.5 49.5 37.1 37.4 46.9 38.7 Th/La 0.23 0.43 0.23 0.40 0.26 0.25 0.36 Ta/Hf 0.142 0.125 0.128 0.116 0.131 0.108 0.116 Th/Hf 1.46 1.54 1.46 1.71 1.29 1.28 1.60 Ba/Cs 288 81 127 110 802 514 117 K/Cs 4630 1396 19033 1934 Th/U 2.8 2.7 2.4 4.1 3.8 3.8 3.6 Latitude °S 37°19.91 37°19.91 37°19.91 37°25.15 37°25.15 37°25.15 37°27.43 Longitude °W 70°26.62 70°23.69 70°23.69 70°23.69 70°23.18 70°23.18 70°23.18 Elevation (m) 1195 1807 1807 1807 872 872 872 Age (Ma) Note: Analytical techniques for major- and trace-element analyses are described in Kay et al. (this volume, chapter 10).
La Ce Nd Sm Eu Tb Yb Lu Sr Ba Cs U Th Hf Ta Sc Cr Ni Co 3.4 1.6 5.4 0.87 28.2 1372 48.6 0.31 0.106 1.59 564 3.4 37°25.15 70°23.18 872
15.5 35.2 20.6 4.62 1.30 0.751 2.85 0.380 619 437 0.8 1.4 4.8 3.0 0.32 19.9 5 8 20
5.99 2.81 3.1 1.4 4.4 0.90 33.6 1384 41.2 0.21 0.102 0.88 323 13143 3.8 37°04.69 70°17.91 1606
10.6 25.9 15.9 3.46 1.09 0.575 2.44 0.332 458 358 1.1 0.6 2.2 2.5 0.26 9.2 3 3 11
5.79 2.36 4.0 1.6 6.5 0.90 35.0 1278 36.5 0.27 0.108 1.05 134 3262 3.5 37°05.99 70°21.53 1530 11.7 ± 0.20
15.3 35.6 16.8 3.81 1.00 0.558 2.35 0.334 462 535 4.0 1.2 4.1 3.9 0.42 9.0 4 3 10
TABLE 3. WHOLE-ROCK CHEMISTRY OF LATE CRETACEOUS TO MIOCENE MAGMATIC ARC ROCKS FROM THE NEUQUÉN ANDES (continued) 16 17 18 19 20 21 22 23 24 25 Fm./Group Late Paleocene to Eocene east of the Cordillera del Viento in Neuquén Miocene Locality Cerro Caicayen Cerro Mayal south Sierra del Mayal Cerro Negro Sample TDR22 TDR28a TDR28d TDR28b TDR23a TDR23b TDR23d TDR23c TDR34 TDR21 SiO2 59.93 58.86 52.62 52.94 60.25 60.77 TiO2 0.53 0.53 1.22 2.04 0.65 0.67 22.25 19.24 19.33 17.18 16.98 17.16 Al2O3 FeO 3.80 4.19 4.37 3.06 7.67 8.14 10.29 8.69 6.78 5.87 MnO 0.16 0.23 0.30 0.18 0.20 0.18 MgO 1.44 3.15 4.22 3.24 2.41 2.49 CaO 6.99 5.67 9.17 8.00 6.57 6.56 4.28 7.39 6.93 3.71 3.33 2.61 3.58 3.87 4.24 4.23 Na2O 1.02 0.94 1.36 1.26 1.75 1.57 K2O 0.22 0.03 0.33 0.49 0.20 0.24 P2O5 Total 100.62 100.24 99.55 99.21 100.02 99.73
Upper Cretaceous to Holocene magmatism 27
VL-3 47.91 1.33 18.53 9.98 0.17 6.59 10.78 2.50 0.29 0.21 98.29
15.5 30.8 19.2 4.61 1.279 0.723 2.18 0.331 491 276 0.2 0.5 2.3 2.6 0.37 35.2 196 80 42
2.79 1.51 3.4 2.1 7.1 0.51 17.8 750 42.1 0.15 0.14 0.90 1673 14580 4.3 36.852 71.067 1687 ca. 24
Fm./Group Sample SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O P2O5 Total
La Ce Nd Sm Eu Tb Yb Lu Sr Ba Cs U Th Hf Ta Sc Cr Ni Co
Na2O+K2O FeO/MgO La/Sm Sm/Yb La/Yb Eu/Eu* Ba/La Ba/Ta La/Ta Th/La Ta/Hf Th/Hf Ba/Cs K/Cs Th/U Latitude °S Longitude °W Elevation (m) Age (Ma)
1
3.05 1.47 3.3 2.0 6.6 0.51 21.2 757 35.7 0.13 0.15 0.68 404 4898 3.0 37.016 71.013 1480 ca. 24
15.2 30.9 19.4 4.57 1.303 0.740 2.29 0.327 483 322 0.8 0.6 1.9 2.8 0.43 30.4 157 69 36
RB-10 48.65 1.40 18.01 9.09 0.20 6.18 10.81 2.58 0.47 0.30 97.69
2
3.02 1.40 2.9 1.7 5.0 0.51 19.5 1014 52.1 0.11 0.10 0.58 1583 27405 4.7 36.829 71.095 1460 ca. 24
10.1 24.0 15.1 3.47 1.151 0.648 2.02 0.280 461 196 0.1 0.2 1.1 2.0 0.19 31.6 118 61 39
LE-5 48.98 1.08 17.84 8.85 0.16 6.33 10.54 2.61 0.41 0.11 96.91
3
3.90 1.54 4.0 2.2 8.6 0.51 21.2 834 39.4 0.14 0.18 1.00 3543 53801 3.6 36.845 71.066 1875 ca. 24
22.4 46.8 26.4 5.62 1.538 0.717 2.61 0.340 621 474 0.1 0.9 3.2 3.2 0.57 30.2 126 57 35 5.97 2.93 2.5 1.7 4.3 0.51 34.8 1434 41.2 0.14 0.11 0.62 1248 29840 3.1 37.025 71.023 1462 ca. 24
13.5 29.9 22.0 5.39 1.577 0.929 3.12 0.461 694 469 0.4 0.6 1.8 3.0 0.33 23.8 7 8 19 7.91 7.51 5.3 1.8 9.6 0.51 31.3 1284 41.0 0.24 0.18 1.75 180 5277 2.9 37.371 70.974 1443 22.8 ± 0.7
29.0 64.9 25.7 5.51 1.106 0.858 3.03 0.444 249 907 5.0 2.4 7.1 4.0 0.71 5.7 3 0 4 9.89 3.75 4.0 2.5 10.1 0.51 17.1 1202 70.5 0.22 0.09 1.42 388 43046 3.7 37.371 70.974 1443 26.3 ± 1.5
17.0 37.9 20.0 4.23 1.227 0.546 1.69 0.235 654 290 0.7 1.0 3.7 2.6 0.24 21.3 44 35 24 4.39 2.91 2.5 1.8 4.5 0.51 14.7 642 43.8 0.18 0.09 0.71 620 20779 3.0 37.223 70.917 1700 ca. 20
20.3 48.8 35.6 8.20 2.064 1.375 4.49 0.646 406 297 0.5 1.2 3.7 5.1 0.46 31.6 7 12 27
ca. 20
4.95 2.20 2.6 1.8 4.6 0.51 18.6 883 47.4 0.15 0.09 0.65 447 14998 3.0 37.214 70.927
11.8 28.3 17.3 4.55 1.340 0.724 2.60 0.374 467 220 0.5 0.6 1.8 2.8 0.25 22.1 8 7 24 4.37 2.62 3.0 2.0 6.1 0.51 17.7 775 43.8 0.28 0.09 1.12 277 9521 3.2 37.374 70.975 1425 ca. 20
17.1 38.2 22.9 5.66 1.354 0.976 2.78 0.414 515 302 1.1 1.5 4.7 4.2 0.39 27.8 6 5 23 3.56 2.35 2.4 1.6 3.8 0.51 21.7 1115 51.4 0.17 0.08 0.69 213 3137 2.8 37.146 70.993 1991 ca. 20
8.6 20.5 13.6 3.54 1.038 0.624 2.26 0.314 416 188 0.9 0.5 1.5 2.1 0.17 25.3 16 7 23 5.75 2.57 3.1 1.9 6.0 0.51 20.9 856 41.0 0.26 0.11 1.10 188 5725 3.3 36.864 70.965 1456 ca. 12
20.8 49.6 29.3 6.74 1.583 1.027 3.47 0.506 502 433 2.3 1.6 5.3 4.8 0.51 23.6 7 7 22
TABLE 4. WHOLE-ROCK CHEMISTRY: MIOCENE CURA MALLIN AND TRAPA TRAPA FORMATIONS 4 5 6 7 8 9 10 11 12 Cura Mallin Fm. Trapa Trapa Fm. ca. 20 Ma LE-3 RB-11 RR6 AL2 RLL-15 RLL-13 RR-5 AP-3 RR-10 49.08 53.42 71.59 67.31 50.61 52.19 51.68 51.74 54.42 1.35 1.57 0.29 0.41 1.89 1.21 1.39 1.16 1.25 18.46 19.97 14.39 15.35 16.10 18.37 18.73 18.79 18.94 9.03 8.94 2.82 3.20 11.31 9.93 9.01 8.94 8.38 0.18 0.19 0.07 0.13 0.23 0.21 0.14 0.28 0.19 5.88 3.05 0.38 0.85 3.89 4.50 3.44 3.80 3.26 10.00 5.12 2.15 2.48 8.30 8.20 9.56 9.22 7.30 3.03 4.62 4.70 6.01 3.19 4.06 3.12 3.22 4.16 0.87 1.35 3.21 3.88 1.20 0.89 1.25 0.33 1.59 0.26 0.27 0.09 0.10 0.58 0.22 0.05 0.40 98.15 98.51 99.70 99.71 97.30 99.55 98.54 97.54 99.88
4.83 2.40 3.6 2.5 9.1 0.51 27.1 1455 53.7 0.19 0.09 0.92 626 14787 2.8 36.864 70.965 1456 ca. 12
13.7 26.3 17.9 3.81 0.935 0.501 1.50 0.225 634 371 0.6 1.0 2.7 2.9 0.25 15.7 3 8 21
4.53 2.51 3.1 1.7 5.2 0.51 23.8 913 38.4 0.34 0.09 1.16 294 9792 3.8 37.223 70.917 1700
15.7 37.4 19.1 4.98 1.228 0.829 3.01 0.432 451 373 1.3 1.4 5.3 4.5 0.41 17.5 3 5 15
6.78 2.18 4.6 1.7 7.9 0.51 24.6 1015 41.3 0.23 0.11 1.07 455 10013 3.5 37.368 71.019 1545 ca. 15
23.3 51.6 23.7 5.03 1.402 0.691 2.94 0.461 508 573 1.3 1.5 5.3 4.9 0.56 11.7 7 3 11
13 14 15 Trapa Trapa Fm. <16 Ma RN-4 RLL-16 RR-9 54.75 62.24 60.20 0.92 0.75 0.71 19.59 16.03 17.04 7.10 6.62 5.83 0.13 0.11 0.18 2.96 2.64 2.67 7.95 6.73 6.41 3.77 3.03 5.26 1.06 1.50 1.52 0.09 0.28 98.32 99.66 100.11
5.63 2.11 3.2 1.4 4.6 0.51 37.1 1327 35.8 0.40 0.08 1.21 140 5293 5.3 36.895 71.114 1642 ca. 15
12.2 27.9 12.3 3.80 0.897 0.649 2.63 0.381 302 452 3.2 0.9 4.9 4.0 0.34 15.1 7 4 14
VL-4 67.29 0.69 13.60 5.46 0.10 2.59 4.54 3.57 2.06 0.14 100.05
16
28 S.M. Kay et al.
7.47 1.77 8.59 Na20+K2O FeO/MgO 3.24 1.97 6.15 La/Sm 4.0 6.7 6.0 Sm/Yb 1.8 0.9 1.4 La/Yb 7.5 5.9 8.4 Eu/Eu* 0.82 0.82 0.82 Ba/La 6.5 2.2 28.7 Ba/Ta 273 82 861 La/Ta 41.8 37.2 30.1 Th/La 0.23 0.04 0.55 Ta/Hf 0.12 1.34 0.22 Th/Hf 1.17 2.00 3.71 Ba/Cs 340 122 106 K/Cs 41766 1575 9641 Th/U 4.1 0.1 3.3 Latitude °S 36°50.34 36°50.14 36°49.50 Longitude °W 70°40.3 70°41.8 70°42.52 Elevation (m) 1322 1397 1180 Age (Ma) Miocene
8.1 15.6 11.9 3.1 0.9 0.6 1.8 0.3 400 198 0.4 0.6 1.7 1.9 0.13 34.8 71 28 28
11.8 25.8 14.3 3.50 1.007 0.492 1.50 0.217 583 259 0.8 0.4 1.7 2.4 0.21 15.9 22 19 24
13.2 24.9 15.1 3.77 1.170 0.631 2.14 0.308 416 270 0.3 0.8 2.4 2.6 0.21 19.0 66 33 20
6.80 3.64 4.77 4.14 3.50 1.72 1.93 1.71 3.3 2.6 3.4 3.5 2.8 1.7 2.3 1.8 9.3 4.4 7.8 6.1 0.82 0.82 0.82 0.82 17.0 24.3 22.0 20.5 707 1536 1238 1305 41.5 63.2 56.3 63.7 0.24 0.21 0.14 0.18 0.09 0.07 0.09 0.08 0.95 0.89 0.71 0.93 348 472 322 903 11978 13137 8998 24463 3.2 3.0 3.9 3.1 36°45.62 36.969° 36.553° 37.008° 70°43.62 70.788° 70.754° 70.816° 1567 1561 1442 1335 7 10? 9
31.5 74.7 45.3 9.48 1.949 1.222 3.40 0.465 466 537 1.5 2.4 7.7 8.0 0.76 24.3 6 7 19
51.1 92.9 52.1 10.83 1.483 1.274 4.14 0.521 195 783 3.8 6.1 20.0 9.8 1.18 10.7 4 2 5
8.85 6.70 4.6 4.7 2.1 2.6 9.6 12.4 0.82 0.82 22.7 15.3 1146 662 50.5 43.2 0.33 0.39 0.11 0.12 1.80 2.03 557 206 10445 3.2 3.3 36.996° 36°45.68 70.851° 70°43.4 1250 1673 10?
23.4 49.9 23.8 5.13 1.159 0.582 2.43 0.342 499 531 1.0 2.4 7.8 4.4 0.46 9.8 5 6 10 4.66 1.83 3.0 2.3 7.0 0.82 21.9 844 38.6 0.20 0.11 0.86 229 8145 3.3 37.343° 70.768° 2.5
14.1 35.2 17.3 4.66 1.312 0.674 2.02 0.292 581 308 1.3 0.9 2.8 3.3 0.36 27.1 66 28 24 4.86 1.94 3.1 2.4 7.3 0.82 21.8 1024 46.9 0.18 0.10 0.82 207 5843 3.2 37.344° 70.886° 1259 2.5
15.3 35.0 22.4 4.96 1.346 0.740 2.10 0.303 611 334 1.6 0.9 2.8 3.4 0.33 27.5 73 33 25 4.37 1.83 3.3 2.2 7.1 0.82 17.4 842 48.3 0.21 0.10 1.01 224 7967 3.4 37.205° 71.144° 1952 2.5
15.4 34.7 21.8 4.66 1.175 0.704 2.16 0.309 430 269 1.2 1.0 3.3 3.2 0.32 22.1 31 26 29
16.5 40.3 20.2 5.22 1.302 0.771 2.36 0.351 483 371 1.6 1.2 4.6 4.5 0.38 27.0 52 21 22
17.9 43.7 23.8 5.58 1.429 0.771 2.61 0.354 504 360 2.5 1.5 5.1 4.8 0.40 25.6 36 21 24
20.7 48.5 24.4 5.85 1.341 0.828 3.28 0.457 410 476 3.1 2.6 8.3 5.1 0.52 27.3 43 13 25
28.0 66.5 38.0 8.24 1.956 1.147 3.67 0.511 474 599 0.6 1.9 7.0 6.6 0.67 18.9 7 3 10 5.12 5.55 2.33 1.96 3.2 3.2 3.2 3.5 3.4 1.8 2.2 2.1 1.8 2.2 5.9 7.0 6.9 6.3 7.6 0.82 0.82 0.82 0.82 0.82 24.2 22.4 20.0 23.0 21.4 1105 985 889 921 895 45.7 43.9 44.4 40.0 41.8 0.23 0.28 0.29 0.40 0.25 0.10 0.08 0.08 0.10 0.10 1.00 1.03 1.08 1.64 1.06 220 237 146 154 1084 6800 7840 2.9 4.0 3.5 3.2 3.6 37.215° 37.220° 37.215° 37.218° 37.219° 71.117° 71.066° 71.053° 70.752° 70.757° 1961 1622 1688 1000 2.5 2.5 2.5 2.5 2.5
14.7 30.9 20.7 4.53 1.199 0.739 2.48 0.356 526 355 1.6 1.1 3.4 3.4 0.32 23.5 19 11 22
La Ce Nd Sm Eu Tb Yb Lu Sr Ba Cs U Th Hf Ta Sc Cr Ni Co
21.3 44.0 15.4 3.55 0.363 0.492 2.55 0.341 58 611 5.8 3.6 11.8 3.2 0.71 3.0 1 1 1
31.0 70.6 30.3 7.70 1.806 1.217 4.16 0.593 149 202 0.6 1.8 7.2 6.1 0.74 11.9 3 1 3
Fm./Group Locality Sample SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O P2O5 Total
111.7 45.3 63.1 16.57 2.323 2.770 19.02 2.911 267 246 2.0 30.0 4.5 2.2 3.00 13.4 17 10 9
TABLE 5. WHOLE ROCK CHEMISTRY OF NEOGENE MAGMATIC ROCKS FROM THE NEUQUEN ANDES WEST OF THE CORDILLERA DEL VIENTO 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 Q. Honda Miocene Cajón Negro Fm. Fm. Miocene/Pliocene Pichi Neuquén Complex Late Pliocene Cola de Zorro, Centinela Pichi Neuquén Complex Río Reñileuvu Chile Río Lileo BPN10 BPN3a BPN4 BPN15 SPN-1 BPN16 RN-1 RN2 BPN13 RR-12 RR-11 RLLC-2 RLL-3 RLL-4 RLL5 RLL2 RLL1 64.50 66.85 76.91 57.40 50.81 53.27 56.24 67.57 52.43 52.83 52.27 54.58 54.96 0.70 0.62 0.10 1.74 1.06 0.85 0.74 0.74 1.34 1.25 1.31 1.12 1.10 17.46 13.39 12.94 15.66 18.66 20.15 19.07 15.91 18.83 18.57 18.04 18.16 17.46 4.34 5.09 0.65 8.72 8.18 7.78 7.07 5.77 3.67 7.66 7.96 8.09 7.90 7.51 8.11 7.85 6.18 0.10 0.07 0.02 0.16 0.18 0.14 0.18 0.07 0.13 0.15 0.12 0.15 0.13 1.34 2.58 0.11 2.49 4.75 4.04 4.12 0.55 4.18 4.11 4.43 3.39 3.84 3.79 6.90 0.64 6.42 10.27 8.91 8.08 2.23 10.00 9.70 8.42 7.65 9.01 4.48 1.39 1.90 4.57 2.98 3.90 3.26 4.85 4.06 3.34 3.72 3.22 3.80 4.07 4.27 3.79 5.26 2.99 0.38 6.69 2.23 0.66 0.87 0.88 4.79 1.32 1.14 1.15 1.32 1.48 0.28 2.63 0.00 0.60 0.01 0.18 0.24 0.14 0.14 0.11 99.98 99.92 99.95 100.01 97.56 100.09 99.89 99.73 99.23 99.43 97.19 98.14 99.80 -
1
3.74 1.41 2.8 1.6 4.5 0.82 24.7 1149 46.6 0.17 0.10 0.79 342 10051 3.5 37.270° 70.818° 1287 1.2-1.4
9.4 20.9 12.0 3.32 0.967 0.564 2.08 0.289 562 232 0.7 0.5 1.6 2.0 0.20 36.0 116 26 30
Quaternary Guañacos RG-9 52.06 0.88 18.16 8.08 0.15 5.72 11.14 2.92 0.82 99.93
18
<1
3.89 1.10 2.7 1.9 5.0 0.82 23.3 1101 47.2 0.16 0.09 0.67 224 5803 4.0
8.8 21.5 13.1 3.31 0.970 0.542 1.76 0.237 516 206 0.9 0.4 1.4 2.1 0.19 25.2 130 106 41
SVZ Antuco Ant1 51.21 1.03 16.87 8.54 0.16 7.73 8.75 3.25 0.64 0.14 98.33
19
Upper Cretaceous to Holocene magmatism 29
Na2O+K2O FeO/MgO La/Sm Sm/Yb La/Yb Eu/Eu* Ba/La Ba/Ta La/Ta Th/La Ta/Hf Th/Hf Ba/Cs K/Cs Th/U Latitude °S Longitude °W Elevation (m) Age (Ma)
4.5 ± 0.5
5.84 2.38 3.6 2.1 7.6 0.90 13.8 192 13.9 0.09 0.36 0.48 636 28696 2.6 37°41.8 69°34'
5.05 1.67 3.7 2.9 10.5 0.90 36.6 570 15.6 0.13 0.43 0.85 503 9008 3.3 37°50.3 69°30.7'
5.82 2.17 4.2 2.3 9.7 0.90 34.9 821 23.5 0.25 0.21 1.22 569 11153 2.9 37°13.33 70°14.34 1443 ± 40
6.38 1.94 4.2 2.8 12.1 0.77 35.0 1032 29.5 0.28 0.20 1.62 203 4596 2.3 37°19.26 70°9.67 1404 ± 39
25.7 54.5 32.9 6.06 1.423 0.817 2.13 0.291 788 897 4.4 3.0 7.1 4.4 0.87 14.0 38 22 19
2.9 37°18.77 70°8.41 1709 ± 32
6.30 2.23 5.0 2.6 13.0 0.91 28.7 962 33.6 0.30 0.17 1.71 274
21.8 46.4 21.1 4.35 1.137 0.518 1.68 0.226 567 625 2.3 2.2 6.4 3.8 0.65 9.4 4 3 15 6.42 2.41 4.1 2.7 11.2 0.79 38.5 1007 26.2 0.28 0.21 1.51 236 5235 2.5 37°18.77 70°8.41 1709 ± 32
23.2 50.8 27.4 5.65 1.397 0.833 2.08 0.315 729 894 3.8 2.6 6.4 4.2 0.89 16.2 29 20 21
23.4 51.7 18.9 3.73 0.945 0.487 1.56 0.194 445 773 3.2 3.4 9.3 4.5 1.01 5.5 4 3 8
6.96 2.21 4.3 6.3 2.3 2.4 9.9 15.0 0.77 0.77 32.5 33.1 769 763 23.7 23.1 0.30 0.40 0.22 0.23 1.52 2.08 280 244 7261 2.8 2.7 37°18.33 ~37°22 70°7.94 ~70°9' 1988 ± 42
23.9 51.1 26.8 5.52 1.319 0.790 2.42 0.341 508 774 2.8 2.5 7.1 4.7 1.01 14.9 18 9 17
37.9 69.8 33.9 6.21 1.299 0.789 2.33 0.276 447 692 2.7 4.6 11.8 5.5 1.22 8.5 6 5 12 6.1 2.7 16.3 0.62 18.3 567 31.0 0.31 0.22 2.14 256 2.6 37°10.1 69°55.9 9.00 9.76 6.8 1.7 11.4 0.19 22.9 268 11.7 0.75 0.44 3.86 116 9276 3.0 37°18.33 70°7.94 1988 ± 42
21.1 42.3 15.1 3.11 0.172 0.395 1.86 0.231 28 484 4.2 5.3 15.8 4.1 1.81 1.9 <1 <1 <1 7.4 1.4 10.7 0.21 20.9 255 12.2 0.73 0.42 3.76 89 2.9 37°18.33 70°7.94 1988 ± 42
22.1 45.0 8.2 2.99 0.173 0.313 2.06 0.282 29 463 5.2 5.5 16.2 4.3 1.82 1.9 1 <1 <1 7.7 1.6 12.2 0.21 18.0 257 14.3 0.59 0.44 3.71 139 2.9 37°18.33 70°7.94 1988 ± 42
26.9 51.7 17.7 3.49 0.167 0.405 2.20 0.277 33 483 3.5 5.5 15.9 4.3 1.88 2.0 1 <1 <1
29.0 57.8 19.7 3.76 0.481 0.155 1.97 0.273 84 733 5.2 5.3 15.2 4.5 1.81 1.9 <1 <1 <1
22.0 43.0 8.5 2.62 0.366 0.184 1.37 0.189 182 632 5.9 5.5 16.8 4.4 1.61 1.9 1 <1 2
4.0 ± 0.4
8.89 9.03 6.82 6.41 7.8 7.7 8.4 2.4 1.9 1.9 18.6 14.7 16.1 0.60 0.62 0.62 24.8 25.3 28.7 427 405 393 17.2 16.0 13.7 0.51 0.53 0.76 0.41 0.41 0.37 3.62 3.42 3.84 148 141 107 6666 6676 3.0 2.9 3.1 37°10.6'' 37°10.1' 37°10.1' 69°56 69°55.9 69°55.9
29.8 59.0 20.8 3.80 0.488 0.175 1.60 0.241 84 737 5.0 5.1 15.2 4.2 1.73 1.8 <1 <1 <1
La Ce Nd Sm Eu Tb Yb Lu Sr Ba Cs U Th Hf Ta Sc Cr Ni Co
20.4 43.2 22.4 4.90 1.355 0.674 2.10 0.278 749 713 1.3 1.7 5.0 4.1 0.87 17.3 8 12 23
21.2 46.0 26.4 5.82 1.852 0.951 2.80 0.399 730 292 0.5 0.8 2.0 4.2 1.52 21.3 6 5 36
Fm./ Group Locality Sample SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O P2O5 Total
26.3 57.6 33.7 7.18 2.107 1.013 2.51 0.388 1001 965 1.9 1.0 3.4 4.0 1.69 25.4 134 47 40
TABLE 6. WHOLE ROCK CHEMISTRY OF PLIOCENE TO HOLOCENE MAGMATIC ROCKS FROM THE TROMEN MASSIF AND CHOS MALAL TROUGH REGIONS 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 Parva Negra Coyocho Formation (Basalt II) Tilhué Fm Rhyolites Cerro Tilhué agglomerate Cerro Bayo Cerro Tilhué Cerro Bayo HDR12 HDR5 TDR35 TDR26 TDR27c TDR27d TDR24c TDR25 TDR17b TDR24a TDR24b TDR24e TDR16 TDR17c TDR17a 47.71 46.07 53.26 56.61 56.79 56.52 59.44 76.01 75.05 75.66 2.32 2.25 1.28 0.92 0.88 1.06 1.11 0.13 0.18 0.17 18.49 17.28 19.55 17.05 16.91 18.29 17.08 13.06 13.96 13.47 11.28 10.58 7.91 6.68 6.33 7.14 6.55 3.96 5.11 1.23 1.03 1.06 1.11 1.09 1.55 0.20 0.02 0.15 0.12 0.14 0.11 0.14 0.08 0.13 0.05 4.75 6.34 3.65 3.45 2.84 2.96 2.97 0.13 0.16 0.17 8.69 10.76 8.08 7.50 9.08 7.25 5.93 0.41 0.71 0.60 4.26 2.97 4.14 3.92 4.16 4.03 4.54 4.49 4.18 4.34 4.34 4.35 4.90 4.85 4.14 1.58 2.08 1.68 2.45 2.14 2.39 2.42 4.66 3.99 4.18 0.49 0.74 0.38 0.43 0.34 0.38 0.31 0.00 0.00 0.00 99.77 99.08 100.09 99.12 99.60 100.14 100.49 100.05 100.19 100.23 -
(continued)
8.56 9.40 5.3 1.5 8.0 0.28 40.2 460 11.5 0.65 0.47 3.52 158 9385 2.3 37°9.64 70°15.09 1561 ± 35
15.7 34.0 12.8 2.94 0.249 0.403 1.96 0.251 38 630 4.0 4.4 10.2 2.9 1.37 1.3 1 <1 <1
pumice TDR20 76.85 0.06 13.12 0.80 0.09 0.09 0.45 4.06 4.50 0.00 100.02
16
30 S.M. Kay et al.
Na2O+K2O FeO/MgO La/Sm Sm/Yb La/Yb Eu/Eu* Ba/La Ba/Ta La/Ta Th/La Ta/Hf Th/Hf Ba/Cs K/Cs Th/U Latitude °S Longitude °W Elevation (m) Age (Ma)
La Ce Nd Sm Eu Tb Yb Lu Sr Ba Cs U Th Hf Ta Sc Cr Ni Co
Fm./group Locality Sample SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O P2O5 Total
24.7 52.0 30.3 6.47 1.544 0.981 2.87 0.396 545 499 0.3 1.4 4.5 4.9 0.73 21.9 20 16 28
26.8 57.7 33.0 6.40 1.676 0.916 2.75 0.368 597 560 1.8 1.9 6.1 5.6 0.90 21.4 17 11 23
5.33 6.03 6.14 2.01 2.70 2.54 4.0 3.8 4.2 2.6 2.3 2.3 10.2 8.6 9.8 0.78 0.76 0.84 17.3 20.2 20.9 471 688 624 27.3 34.1 29.8 0.25 0.18 0.23 0.17 0.15 0.16 1.17 0.93 1.10 495 1518 318 15022 36621 9216 3.6 3.3 3.3 37°11.09 37°08 37°5.99 70°8.83 70°10.95 70°09 2158 ± 30 1912 ± 56
23.5 52.8 22.9 5.91 1.443 0.885 2.31 0.318 654 406 0.8 1.7 6.0 5.1 0.86 19.5 38 26 27
21.1 44.8 24.6 5.79 1.490 0.869 2.76 0.377 464 616 1.3 1.3 4.5 4.7 0.66 28.7 42 21 32
4.71 5.20 2.35 2.04 3.5 3.7 2.1 2.1 7.3 7.6 0.83 0.82 24.3 29.2 660 936 27.2 32.1 0.18 0.21 0.18 0.14 0.87 0.96 706 492 13265 9052 3.2 3.4 37°10.77 37°02.78 70°14.95 70°19.70 1412 ± 62 1336 ± 30 1.44 ± 0.08
21.8 45.3 26.8 6.20 1.620 0.926 3.00 0.407 452 530 0.8 1.2 4.0 4.6 0.80 28.9 32 25 31 5.38 1.98 3.8 2.2 8.4 0.88 16.4 432 26.3 0.24 0.24 1.53 671 25282 3.4 37°03.50 70°01.8
21.1 45.6 21.4 5.50 1.533 0.844 2.53 0.357 607 346 0.5 1.5 5.1 3.3 0.80 20.4 76 45 27 6.18 2.76 4.1 2.3 9.3 0.87 19.9 467 23.5 0.22 0.22 1.15 380 10053 3.2 37°03.4 70°08.4
26.9 59.5 33.3 6.61 1.803 0.977 2.89 0.399 534 534 1.4 1.9 6.0 5.2 1.14 23.5 4 4 23 6.13 2.42 4.2 2.3 9.8 0.83 20.0 597 29.8 0.23 0.16 1.11 257 7682 3.3 37°10.20 70°9.21 2202 ± 41
26.9 57.7 32.9 6.39 1.656 0.917 2.74 0.368 614 539 2.1 1.9 6.3 5.7 0.90 21.4 21 13 24
1.04 ± 0.06
5.44 1.48 5.1 3.1 15.5 0.90 20.9 647 31.0 0.30 0.18 1.69 209 5095 3.4 37°06.50 70°05.06
24.2 51.9 20.9 4.77 1.253 0.573 1.56 0.224 710 504 2.4 2.1 7.2 4.3 0.78 17.6 95 57 26 8.21 3.53 4.8 2.2 10.5 0.69 19.4 547 28.3 0.33 0.15 1.36 174 6501 3.3 37°03.6 70°09.5
38.7 85.6 36.8 8.13 1.691 1.066 3.67 0.527 344 749 4.3 3.9 12.7 9.3 1.37 11.3 4 2 6 4.34 1.36 3.0 2.0 6.0 0.88 16.6 327 19.8 0.27 0.19 1.04 200 7694 3.1 37°03.0 69°57.60
12.1 27.4 15.1 3.98 1.154 0.684 2.00 0.287 380 200 1.0 1.0 3.3 3.2 0.61 28.3 206 91 44 6.96 1.94 5.4 2.5 13.3 0.67 17.8 377 21.2 0.49 0.20 2.12 102 4109 3.3 37°03.5 69°55.73
32.2 68.8 29.2 6.01 1.230 0.827 2.43 0.285 410 573 5.6 4.8 15.8 7.5 1.52 14.4 27 15 18
6.61 2.09 5.3 2.6 14.0 0.81 22.2 510 23.0 0.42 0.20 1.94 158 5599 3.1 37°06.50 70°05.06
26.8 55.2 23.8 5.01 1.176 0.577 1.91 0.264 533 595 3.8 3.6 11.2 5.8 1.17 14.1 22 17 18
0.18 ± 0.04
7.41 2.16 5.3 2.5 12.9 0.68 20.8 431 20.7 0.56 0.19 2.23 113 4647 3.2 37°05.50 70°05.5
30.3 63.4 26.3 5.75 1.162 0.736 2.35 0.359 492 631 5.6 5.4 17.1 7.7 1.47 13.0 24 15 17
7.90 2.38 4.5 2.4 11.0 0.85 21.1 372 17.6 0.39 0.25 1.77 95 4273 2.6 37°09.15 69°48.9
29.4 61.9 31.0 6.48 1.517 0.658 2.68 0.367 599 622 6.5 4.4 11.6 6.6 1.67 8.9 2 2 8
7.67 2.99 4.2 2.4 10.0 0.76 18.1 310 17.1 0.41 0.25 1.73 111 5770 2.7 37°09.71 69°57.5
29.0 60.9 31.3 6.84 1.515 0.813 2.91 0.392 699 526 4.7 4.3 11.9 6.9 1.69 8.2 4 5 10
TABLE 6. WHOLE ROCK CHEMISTRY OF PLIOCENE TO HOLOCENE MAGMATIC ROCKS FROM THE TROMEN MASSIF AND CHOS MALAL TROUGH REGIONS (continued) 17 18 19 20 21 22 23 24 25 26 27 28 29 30 31 32 "Chos Malal Group" Tromen (Basalts VI, VII) Agua Carmonina Chapúa Fm. Maipo Fm. El Puente Fm. C° Waile Michico escorial' agglomerate blocks TDR31 TDR29 TDR7 TDR19 TDR33 TDR1 TDR5 TDR30 TDR3 TDR6 TDR12 TDR13 TDR4 TDR2 TDR15 TDR18 52.25 54.76 54.85 50.11 52.02 51.08 53.99 54.89 54.84 63.54 50.33 59.25 59.49 60.66 58.40 59.30 1.47 1.52 1.53 1.84 1.59 1.23 1.63 1.47 1.08 0.94 1.46 1.21 0.94 1.01 0.86 0.91 19.77 17.44 18.30 18.00 16.86 19.34 16.94 18.41 18.06 16.56 17.18 16.66 17.42 17.00 19.20 19.63 8.14 9.10 8.08 10.42 9.80 8.40 8.84 8.13 7.01 4.44 9.78 6.18 6.03 6.07 5.07 4.26 0.15 0.16 0.12 0.19 0.18 0.15 0.17 0.14 0.11 0.12 0.15 0.11 0.10 0.18 0.09 4.05 3.37 3.18 4.43 4.80 4.25 3.20 3.36 4.73 1.26 7.21 3.19 2.88 2.81 2.13 1.43 8.50 7.35 7.33 9.35 8.32 9.15 7.12 7.51 7.46 3.26 9.86 5.64 5.70 5.10 5.84 5.13 3.85 4.58 4.19 3.51 3.83 3.81 4.47 4.19 3.96 4.84 3.41 4.17 4.07 4.29 4.54 4.39 1.48 1.45 1.96 1.20 1.37 1.57 1.70 1.94 1.48 3.38 0.93 2.79 2.54 3.12 3.35 3.28 0.40 0.34 0.40 0.42 0.32 0.36 0.45 0.41 0.30 0.30 0.20 0.32 0.26 0.20 0.45 0.33 100.07 100.07 99.92 99.47 99.10 99.34 98.52 100.46 99.03 98.63 100.50 99.41 99.43 100.36 100.03 98.75
Upper Cretaceous to Holocene magmatism 31
32
S.M. Kay et al. TABLE 7. Sr, Nd, AND Pb ISOTOPIC RATIOS FOR NEUQUÉN BASIN REGION MAGMATIC ROCKS Location SiO2 87Sr/86Sr 143Nd/144Nd +eNd 206Pb/204Pb 207Pb/204Pb 208Pb/204Pb Eocene Cayanta Fm. 52.8 0.704546 0.512776 2.7 initial ratios at 52 Ma (32 ppm Rb) 0.704359 3.0 TDR21 Cerro Negro andesite 60.8 0.704744 0.512802 3.2 initial ratios at 11.7 Ma (75 ppm Rb) 0.704665 3.3 TDR24a Cerro Tilhué rhyolite 76.0 0.704358 0.512776 2.7 18.555 15.617 38.476 initial ratios at 4 Ma (100 ppm Rb) 0.704105 2.7 TDR16 Cerro Bayo rhyolite 75.1 0.704407 0.512747 2.1 initial ratios at 1.5 Ma (115 ppm Rb) 0.704211 2.2 TDR19 Chapúa basalt north of Cerro Waile 50.1 0.703851 0.512805 3.2 TDR31 Chapúa basaltic andesite - south of Cerro Waile 52.0 0.704031 0.512789 2.9 18.538 15.599 38.441 TDR6 Cerro Waile andesite 63.5 0.704382 0.512709 1.4 TDR2 Cerro Tromen "escorial" andesite 60.7 0.703927 0.512815 3.4 18.555 15.621 38.472 Standard values NBS987 La Jolla Pb isotopes corrected to NBS981 0.710235 0.511841 16.939 15.497 36.717 Note: Analytical techniques are as in Kay et al. (this volume, chapter 10). Sample ESA12
Nevazón region and 60.7 ± 1.9 Ma from a diorite in the Campana Mahuída region (Franchini et al., 2003). Franchini et al. (2003) showed that these magmatic rocks have chemical signatures similar to those of Holocene Southern Volcanic Zone arc volcanic rocks (see Figs. 5–8). The chemical analyses for the Varvarcó granodiorite in Table 3 shows a similar REE pattern (La/Yb = 6.6) and arc-like La/Ta (28) and Ba/La (31) ratios. Latest Paleocene and Eocene events produced most of the magmatic rocks mapped as the Serie Andesitic and Molle Formations by Groeber (1946) and reassigned to a volcanic Cayanta Formation by Rapela and Llambías (1985) and a plutonic Collipilli Formation by Llambías and Rapela (1989). A distinctive feature of these rocks is the presence of amphibole phenocrysts. The Cayanta Formation is dominantly composed of the extensive breccias flows and necks on the west side of the Cordillera del Viento (Figs. 2A and 3). Their age is confirmed by 40Ar/ 39Ar hornblende ages of 56.0 ± 0.6 Ma and 50.3 ± 0.6 Ma on two flows (Jordan et al., 2001) and wholerock K/Ar ages of 54.2 ± 2.7 Ma on an aplitic stock and 46.1 ± 2.3 Ma on an andesitic dike (Rovere, 1998). The plutonic Collipilli Formation included most of the hornblende-bearing plutonic rocks on the east side of the Cordillera del Indio and was assigned an Eocene age based on K/Ar dates of 49.9 ± 3.2 Ma for the Las Mellizas laccolith in the Collipilli region, 48.4 ± 2.4 Ma at Cerro del Diablo, and 44.7 ± 2.2 Ma at Cerro Caicayén (Llambías and Rapela, 1989). Cobbold and Rossello (2003) reported a whole-rock 40Ar/ 39Ar age of 39.7 ± 0.2 Ma on a sill at Cerro Mayal. Just north of the Rio Barrancas in Mendoza, Linares and González (1990) reported a K/Ar age of 50 ± 5 Ma for hornblende-bearing volcanic rocks in the Cerro Bayo de la Esperanza complex. To avoid confusion with the ages cited above that show that some of the magmatic rocks included in the Collipilli Formation in the Collipilli region are Cretaceous in age (Zamora Valcarce et al., this volume, chapter 6), the magmatic rocks east of the Cordillera del Viento are here informally called the Caicayén group. New chemical analyses of Cayanta Formation, Cerro Bayo de la Esperanza region, and Cerro Mayal region samples listed in Table 3 supplement those from Paleogene samples in Rapela and Llambías (1985), Llambías and Rapela (1989), and Fran-
chini et al. (2003). As seen in Figures 5–8, all of these samples have arc-like features indicated by relative HFSE depletion (La/Ta > 28; Ta/Hf < 0.15) and fluid mobile element enrichment (Ba/La > 20). In detail, there are differences among them. The Cayanta Formation samples west of the Cordillera del Viento are generally similar to Holocene Southern Volcanic Zone arc samples. In both cases, basaltic to mafic andesitic samples have high Al and low Ti contents, arc-like La/Ta (40–64), Ba/La (21–27), and Ta/Hf (0.07–0.12) ratios, and relatively flat REE patterns (La/Yb = 4–7, La/Sm = 3.3–4.1, Sm/Yb = 1.5–1.9). The initial 87Sr/ 86Sr and 143Nd/ 144Nd ratios of a Cayanta basaltic andesite (Table 7) are near those of the Southern Volcanic Zone lavas (Fig. 9). The Cayanta Formation flows differ from the Southern Volcanic Zone lavas in that they typically have amphibole phenocrysts. In contrast, Cerro Bayo de La Esperanza and Cerro Caicayén region samples are more like early Paleocene samples in that they have higher alkali contents, higher La/Ta, Ba/Ta, La/Yb, La/Sm, and Sm/Yb ratios, and lower Ta/Hf ratios than Cayanta samples (Figs. 5–8). Mafic Bayo de la Esperanza region samples (48%–57% SiO2) are particularly notable for their high Na2O (4.4%–6.0%), Sr (604–1402 ppm), and Ba (to 1835 ppm) contents and high La/Ta (64–72), La/Sm (up to 6.7), and Sm/Yb (2.9–4.5) ratios. Cerro Caicayén quartz diorites (59%–62% SiO2) also have high La/Ta* (55–100, where Ta* = Nb/16), Ba/Ta* (most >1400), La/Yb (>9) and La/Sm (up to 14) ratios, but differ in having lower Sm/Yb ratios (<2.3) (Franchini et al., 2003). Cerro Mayal region samples have alkali contents and La/Ta (29–50), Th/Hf (1.3–1.6), La/Yb (5–9), La/Sm (2.6–4.3), and Sm/Yb (1.6–2.2) ratios more similar to the Cayanta samples, but differ in having higher Ta/Hf ratios. Chemical differences between Cerro Caicayén and Cerro Mayal samples could be temporal given the 39.7 ± 0.2 Ma age on the Cerro Mayal sill (Cobbold and Rossello, 2003). Miocene Magmatic Rocks West of the Cordillera del Viento Latest Oligocene to Miocene arc and backarc magmatism in the Neuquén Basin began after a regional magmatic hiatus that lasted from ca. 39 Ma to ca. 26 Ma. The oldest magmatic
Figure 5. Weight percent (K2O + Na2O) versus SiO2 for Neuquén Basin region magmatic samples relative to alkali classification diagram for igneous rocks. Analyses are from Tables 3–6 (points indicated), as well as Franchini et al. (2003) for Paleocene samples; Nullo et al. (2002) and Baldauf (1997) for Huincán I and II, Palaoco, and Molle samples; Kay and Copeland (this volume, chapter 9) for early Miocene backarc samples; Kay et al. (this volume, chapter 10) for Chachahuén and Vizcachas samples; Burns (2002) for Cura Mallín, Trapa Trapa, and Cola de Zorro samples; Kay (2001b) for Auca Mahuída and Río Colorado samples; and Tormey et al. (1991), and references within, for Southern Volcanic Zone (SVZ) samples.32
Figure 6. La/Ta ratios versus (A) SiO2 (wt%) and (B) Ba/Ta ratios for Neuquén Basin region magmatic samples. Shaded region in part B shows general range of ratios in arc magmas. Data sources are as in Figure 5. SVZ—Southern Volcanic Zone.
34 S.M. Kay et al.
Figure 6 (continued).
Figure 7. Th/Hf versus Ta/Hf ratios for Neuquén Basin region magmatic samples. Th/Hf ratios are a relative measure of source components from a subducting oceanic slab, and Ta/Hf ratios are a relative measure of mantle enrichment. Ta and Hf are both high field strength elements; Ta is sensitive to oxidation as it replaces Ti, whereas Hf does not. General ranges of values commonly observed in intraplate, backarc, and arc settings are indicated. Data sources are as in Figure 5. SVZ—Southern Volcanic Zone.
Figure 8 (on this and following page). Rare earth element (REE) characteristics of Neuquén Basin region samples: (A) La/Yb ratio versus SiO2 (wt%) and (B) La/Sm versus Sm/Yb ratios. La/Yb ratios indicate overall REE slopes; La/Sm and Sm/Yb ratios indicate light and heavy REE slopes. High La/Yb and Sm/Yb ratios in basaltic magmas generally indicate small percentage melts of a garnet-bearing mantle source. High ratios in more silicic magmas can indicate that the magmas equilibrated with amphibole-, garnet-, or accessory mineral–bearing residual mineral assemblages. Data sources are as in Figure 5. SVZ—Southern Volcanic Zone.
Figure 8 (continued).
Upper Cretaceous to Holocene magmatism
39
Figure 9. Plot of initial 87Sr/ 86Sr ratios versus εNd values for Neuquén Basin region magmatic samples compared to fields for Holocene Southern Volcanic Zone arc (SVZ; patterned fields). Southern Volcanic Zone data are from compilation of Saal (1994) and Saal et al. (1995) and include data from Hildreth and Moorbath (1988), Stern et al. (1990), and Hickey et al. (1986), among others. Data in field labeled backarc 35°–37°S are from Saal (1994) and Saal et al. (1995). Other Neuquén Basin region data are from Table 7 (points), Kay (2001b), Kay et al. (this volume, chapter 10), and Kay and Copeland (this volume, chapter 9).
rocks of this age in the arc region are in the Cura Mallín Formation that crops out between the Holocene Southern Volcanic Zone arc front and the Cordillera del Viento (Figs. 2A and 3). Most of these magmatic rocks are in sequences of reworked silicic pyroclastic deposits that locally reach thicknesses of ~3 km (Suárez and Emparan, 1995; Burns et al., this volume, chapter 8). These sequences include rare basaltic andesitic lavas and are locally cut by mafic dikes and granitoid intrusives. Jordan et al. (2001) reported hornblende 40Ar/ 39Ar ages of 24.6 ± 1.8 Ma for a basaltic dike in the lower part of the Cura Mallín Formation and of 22.8 ± 0.7 Ma for an ash-flow tuff near the top. Burns et al. (this volume, chapter 8) report a zircon fission-track age of 26.3 ± 1.5 Ma for a granitoid cutting the pyroclastic sequence. The top of the Cura Mallín Formation is generally designated at a change from predominantly sedimentary rocks to volcanic tuffs, agglomerates, and lava flows of Miocene age. The volcanic rocks and related dikes, sills, and granitoid intrusives in Chile are mapped in the Trapa Trapa Formation (Niemeyer and Muñoz, 1983). Their ages are constrained by K/Ar ages that
range from 19.7 ± 1.4 Ma to ca. 12 Ma (Niemeyer and Muñoz, 1983; Muñoz and Niemeyer, 1984; Suárez and Emparan, 1995). Temporally equivalent volcanic rocks in Argentina are assigned to the Trapa Trapa, Cajón Negro, Quebrada Honda, and Pichi Neuquén Formations (see Burns et al., this volume, chapter 8; Figs. 2A and 3). Those in the Trapa Trapa Formation generally occur near the Chilean border. Their ages are constrained by whole-rock K/Ar ages of 18.5 ± 0.2 Ma on a basaltic flow, 12.6 ± 0.2 Ma on an andesite dike, and 12.3 ± 0.2, 12.1 ± 0.6, and 10.8 ± 0.1 Ma on granodiorite and aplitic intrusives (Rovere, 1998). Volcanic rocks assigned to the Cajón Negro Formation, Quebrada Honda Formation, and Pichi Neuquén Volcanic Complex (Pesce, 1981) crop out north of 37°S and east of the Cura Mallín and Trapa Trapa Formation (Fig. 3). They occur north of the westward projection of the Cortaderas lineament (Fig. 2A). The Cajón Negro Formation dominantly consists of mafic andesitic to dacitic flows and agglomerates in the west. To the east are pyroclastic units that include surge deposits, ash falls, tuffs characterized by large pumice fragments, and columnar
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S.M. Kay et al.
Figure 10. 207Pb/ 204Pb and 208Pb/ 204Pb versus 206Pb/ 204Pb for Neuquén Basin region magmatic samples compared to those for Pleistocene to Holocene Southern Volcanic Zone (SVZ) arc samples. Southern Volcanic Zone data are mainly from Hildreth and Moorbath (1988), Stern et al. (1990), and Hickey et al. (1986), among others. Other data are from Table 7 (points), Kay (2001b), Kay et al. (this volume, chapter 10), and Kay and Copeland (this volume, chapter 9). NHRL—northern hemisphere reference line.
Upper Cretaceous to Holocene magmatism jointed ignimbrite flows in various stages of welding. Jordan et al. (2001) reported a hornblende 40Ar/ 39Ar age of 16.2 ± 0.2 Ma from one of the well-preserved silicic tuffs. The composition and preservation of the dated tuff raise the question as to if the dated hornblende could be a xenocryst. Other 40Ar/ 39Ar ages from the Cajón Negro Formation are a biotite age of 11.7 ± 0.3 Ma and a hornblende age of 10.8 ± 1.6 Ma (Burns et al., this volume, chapter 8). The overlying Quebrada Honda Formation dominantly consists of massive gray olivine and pyroxene-bearing columnar jointed andesitic flows. The Pichi Neuquén Complex includes andesitic to dacitic lavas, dikes, necks, and stocks. Jordan et al. (2001) reported a hornblende 40Ar/ 39Ar age of 9 ± 2 Ma for a basal flow. Similarities in style and preservation with volcanic units to the south suggest that part of the Pichi Neuquén Complex could be Pliocene in age, as argued by Pesce (1981). Chemical analyses of Cura Mallín, Trapa Trapa, Cajón Negro, Quebrada Honda, and Pichi Neuquén volcanic rocks and related plutons are listed in Tables 4 and 5 and plotted in Figures 5–8. Overall, these samples are similar to the Eocene Cayanta samples. Among their chemical characteristics are arc-like La/Ta (most 35–55), Ba/La (15–35, most 17–24), and Ta/Hf (most 0.9–0.11) ratios. The Cura Mallín basaltic samples are distinctive in having higher Ta/Hf ratios (0.14–0.18; Table 4). Miocene Magmatic Rocks East of the Cordillera del Viento Miocene magmatic rocks are also abundant east of the Cordillera del Viento (Table 1; Fig. 2A). Between 36.5°S and 38°S, the volcanic rocks in the backarc include early Miocene alkali basaltic flows extending from the middle to the far backarc, subsequent early Miocene volcanic centers in the middle backarc, Miocene volcanic rocks near the east side of the Cordillera del Viento, and the latest Miocene Chachahuén volcanic complex in the far backarc. The oldest of these magmatic rocks are alkali olivine basalts erupted from monogenetic and polygenetic centers in the Sierra de Huantraico (e.g., Ramos and Barbieri, 1988) and the Sierra de Chachahuén and Matancilla region (e.g., González Díaz, 1979). Kay and Copeland (this volume, chapter 9) show that these basalts, which were mapped by previous workers in the Palaoco Formation, have ages from ca. 24 to 20 Ma and are contemporaneous with the Cura Mallín Formation in the arc. Chemical and isotopic data for these basalts from Kay and Copeland (this volume, chapter 9) are plotted in Figures 5–9. As discussed by these authors, their chemical characteristics are those of intraplate basalts generated by partial melting of an isotopically enriched – Nd = +3.6 garnet-bearing mantle (La/Yb ~ 13–30; Sm/Yb > 3.5; C 87 88 to +4.2; Sr/ Sr = 0.7037–0.7040) virtually devoid of arc-like components (La/Ta < 14; Ba/La <16; Ta/Hf >0.45). The least arc-like signatures are found in the Sierra de Chachahuén and Matancilla region magmas, which erupted farthest from the arc (La/Ta <12; Fig. 6). No evidence for contractional deformational structures has been reported for this period. Outcrop-scale extensional faults are present in the basalts in the Sierra de Chachahuén (Ragona, 1999, personal commun.; Kay, 2001b).
41
The ca. 24–20 Ma alkali basalts are succeeded by ca. 19–18 and 16–15 basaltic to hornblende-bearing mafic andesitic to trachydacitic magmas erupted from volcanic complexes (circles in Fig. 2A) in the Sierras de Huantraico and Negra and in southern Mendoza (Nullo et al., 2002; Cobbold and Rossello, 2003; Kay and Copeland, this volume, chapter 9). As shown in Figures 5–9, the chemistry of these younger magmas differs in that they have weak arc-like signatures (La/Ta = 15–26; Ba/La = 15–32; Ta/Hf = 0.2–0.45), flatter REE patterns (Sm/Yb < 3; – Nd = La/Yb = 5–25) and less-enriched isotopic signatures (C 87 88 206 204 +3.9 to +4.7; Sr/ Sr = 0.7033–0.7037; Pb/ Pb = 18.51–18.55). Kay and Copeland (this volume, chapter 9) use these characteristics to argue that these younger magmas were generated in a mantle containing subducted components. They further argue that these magmas erupted in a contractional setting. Evidence for a contractional regime in the Sierra de Huantraico comes from seismic lines that show growth strata associated with folded sequences (Viñes, 1990; see Cobbold and Rossello, 2003) that contain early Miocene volcanic rocks. The next magmas that erupted east of the Cordillera del Viento in Neuquén are the andesites at Cerro Negro (Fig. 2A). A hornblende at that locality yielded the new 40Ar/ 39Ar age of 11.7 ± 0.2 Ma in Table 2. Chemical analyses in Table 3 show that these flows are medium-K andesites (60% SiO2) with arclike La/Ta (37–41), Ba/La (34–35), and Ta/Hf (0.10) ratios and relatively flat REE patterns (La/Yb = 4.4–6.4; La/Sm = 3–4; Sm/Yb = 1.4–1.6). As seen in Figures 5–8, their chemistry is similar to that of Eocene Cerro Mayal lavas to the south and contemporaneous Trapa Trapa andesites in the arc to the west. – Nd of a Cerro Negro andesite are The initial 87Sr/ 86Sr ratio and C in the same general range as those of Eocene to Holocene arc lavas to the west (Fig. 9). The eruptions of the Miocene Trapa Trapa arc volcanic rocks and the Cerro Negro backarc lavas overlap in time with those of the Farellones Formation–Teniente Complex in the arc to the north (e.g., Kay et al., 2005) and the Huincán Formation in the backarc in Mendoza (Fig. 2B; Nullo et al., 2002). The 40Ar/ 39Ar ages reported by Baldauf (1997) for the Huincán Formation between 34°S and 35°S are 13.57 ± 0.9 Ma and 13.94 ± 0.08 Ma for their Huincán I group samples and 10.42 ± 0.05 Ma and five between 7.49 and 5.3 Ma for their Huincán II group samples. Analyses in Baldauf (1997) and Nullo et al. (2002) show that the chemical character of Huincán andesites changed with time. As shown in Figures 5–8, the Huincán II andesites have more arc-like La/Ta (41–61), Ba/La (35–46), and Ta/Hf (0.09–0.13) ratios and steeper REE patterns (La/Yb = 13–17; La/Sm = 4.9–6.5; Sm/Yb = 2.1–2.3) than Huincán I andesites (La/Ta = 34–37; Ba/La = 19–24;Ta/Hf = 0.13–0.17; La/Yb = 9.5–14.5; La/Sm = 4.3–6.4; Sm/Yb = 1.8–3.1). All have steeper REE patterns (higher La/Yb; Fig. 8), than the Cerro Negro andesites to the south. The easternmost Miocene eruptions in the Neuquén Basin occurred at the Chachahuén volcanic complex in the Sierra de Chachahuén (Fig. 2B), some 500 km east of the modern trench
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(Fig. 1). As described by Kay et al. (this volume, chapter 10), the Chachahuén complex consists of the ca. 7.3–6.9 Ma Vizcachas group orthopyroxene-bearing dacites and rhyolites and the ca. 6.8–4.9 Ma high-K Chachahuén group basalts, hornblende-bearing andesites, and dacites. All of these magmas erupted from a nested caldera complex located at the intersection of a NW-SE–trending and a NE-SW–trending fault system. These faults are among those inverted in the late Miocene contractional deformation that uplifted the Sierra de Chachahuén (Pérez and Condat, 1996; Kay et al., this volume, chapter 10). Chemical and isotopic analyses of the Chachahuén complex rocks presented by Kay et al. (this volume, chapter 10) are summarized in Figures 5–9. Chachahuén rocks can be distinguished from those erupted elsewhere in the Neuquén Basin by a combination of high alkali contents (Fig. 5; particularly K), variable La/Ta (10–55; Fig. 6), high Ta/Hf ratios (>1.8, Fig. 7), variable La/Yb, La/Sm, and Sm/Yb ratios (Fig. 8), low 87Sr/ 86Sr ratios at a given εNd (Fig. 9), and high 206Pb/ 204Pb ratios (Fig. 10). Kay et al. (this volume, chapter 10) attribute the intraplate-like character of the older Vizcachas lavas (La/Ta <22; Ta/Hf = 0.4–0.7) to contamination of mantle-generated magmas by intraplatelike crustal melts and the arc-like character of the younger Chachahuén lavas (La/Ta = 32–55; Ta/Hf < 0.3) to melting in the mantle wedge above a subducting slab. The high values of La/Yb (>18), La/Sm (~6–11), and Sm/Yb (>3) of the silicic Vizcachas lavas are attributed to removal of accessory REEbearing phases (e.g., titanite), the importance of which was diminished in the Chachahuén lavas (La/Yb = 7–15; La/Sm ~ 2.5–7; Sm/Yb ~ 1.8–3.2). The Chachahuén volcanic complex is considered to be contemporaneous with the petrologically similar Plateado-Nevado center to the north (Fig. 2B; Bermúdez, 1991; Bermúdez et al., 1993). Pliocene to Holocene Magmatic Rocks West of the Cordillera del Viento Pliocene to Pleistocene volcanic rocks are widespread between the Holocene Southern Volcanic Zone arc and the western side of the Cordillera del Viento (Figs. 2B and 3). Early Pliocene volcanic rocks that form the base and surrounding region of the Pleistocene to Holocene arc are assigned to the Cola de Zola Formation, which has an age of ca. 5.6–3 Ma (Vergara and Muñoz, 1982). Pliocene basaltic to dacitic rocks to the east in Argentina between 37°S and 37.5°S are assigned to the Centinela Formation, K/Ar ages of which range from 3.2 ± 0.2 to 2.6 ± 0.1 Ma (Rovere, 1998). Farther north, a similar age range seems reasonable for the younger part of the Pichi Neuquén Complex. The Pliocene Centinela Formation flows near the western side of the Cordillera del Viento are followed by the Pleistocene olivine basalt flows of the Guañacos Formation, which has K/Ar ages of 1.4 ± 0.2 to 1.2 ± 0.1 Ma (Rovere, 1998). The Cerro Colorado olivine basalt flow mapped on the west side of the Cordillera del Viento to the north by Pesce (1981) is likely to be of this age. The characteristics of other similar late Miocene to Quaternary volcanic rocks northwest,
west, and southwest of the Cordillera del Viento are summarized by Miranda et al. (this volume, chapter 13), Lara and Folguera (this volume, chapter 14), and Folguera et al. (this volume, chapter 12). Holocene Southern Volcanic Zone centers are discussed by Hildreth and Moorbath (1988), Tormey et al. (1991), and Dungan et al. (2001) and references therein. Southern Volcanic Zone rocks from the Copahue volcano near 38°S are addressed by Varekamp et al. (this volume, chapter 15). The chemistry of representative samples from the Pichi Neuquén, Centinela, and Guañacos Formations is presented in Table 5 and plotted in Figures 5–8. These analyses along with those in Rovere (1998) show that these Pliocene to Pleistocene centers have arc-like characteristics similar to those in the Eocene Cayanta Formation, Miocene Cura Mallín and Trapa Trapa Formations, and younger Southern Volcanic Zone centers (Planchon, Cerro Azul, San Pedro, Antuco, Llaima, Villarrica and Puyehue; see Tormey et al., 1991). A new analysis of an Antuco basalt (51% SiO2) in Table 5 with La/Ta = 47, Th/Hf = 0.679, Ta/Hf = 0.09, Ba/La = 23, La/Yb = 5.0, and Sm/Yb = 2.7 is representative of the Southern Volcanic Zone samples. Pliocene to Holocene Magmatic Rocks East of the Cordillera del Viento Pliocene to Holocene arc magmas are contemporaneous with voluminous backarc magmas erupted east of the Cordillera del Viento (Table 1; Fig. 2B). Between 36°S and 38°S, these eruptions produced the small mafic Pliocene flows west of the Auca Mahuída and Payún Matrú fields, the Pliocene to Holocene mafic to silicic volcanic rocks of the Tromen region (Fig. 4), and the voluminous latest Pliocene to Pleistocene mafic alkaline flows in the Auca Mahuída, Payún Matrú, and Llancanelo fields. Early Pliocene Backarc Alkaline Flows. The oldest postMiocene alkaline volcanic rocks in the transect (Fig. 2B) include the flows west of the Payún Matrú field (González Díaz, 1979) and the Parva Negra and Horqueta Norte flows west of the Auca Mahuída field (Ramos and Barbieri, 1988). González Díaz (1979) argued for a Pliocene age for the flows west of the Payún Matrú field on the basis of geomorphology and two whole-rock K/Ar ages of 8 ± 4 Ma and 4 ± 1 Ma. Ramos and Barbieri (1988) reported a K/Ar whole-rock age of 4.5 ± 0.5 Ma for the Parva Negra flow. Chemical analyses for the Parva Negra and Horqueta Norte flows are presented in Table 6 and plotted in Figures 5–8. These flows are olivine alkaline basalts (46%–48% SiO2) with intraplate-like low La/Ta (14–16) and Ta/Hf (0.36– 0.43) ratios and Ba/Ta (192–570) and Ba/La ratios (14–37) that show a variable intraplate signature. The Near to Middle Backarc—Tromen Region. The distribution of post-Miocene volcanic units in the Tromen region is shown on the maps in Figures 2B and 4. The more detailed map in Figure 4 is based on the Chos Malal and Buta Ranquil geologic maps of Zollner and Amos (1973) and Holmberg (1976) and uses the volcanic stratigraphy of Groeber (1946). New 40Ar/ 39Ar ages in Table 2 and locations of samples with analy-
Upper Cretaceous to Holocene magmatism ses in Table 6 are shown on Figure 4. The magmatic sequences on Figure 4 are divided into three groups in the following discussion: (1) an older andesitic and rhyolitic group that includes the Coyocho (Basalto II) and Tilhué (Andesite III) Formations, (2) an intermediate-age basaltic to mafic andesitic group that includes the Chapúa (Basalto III), Maipo (Basalto IV), and El Puente (Basalto V) Formations and is here designated the Chos Malal group, and (3) a younger, mostly andesitic Tromen Formation group that includes Basaltos VI and VIII. The oldest group incorporates the eroded Coyocho Formation mafic flows west of Cerro Waile and Cerro Tilhué, east of Cerro Michico and in the Cerro Bayo region. They are mapped as being older than the Tilhué Formation, which includes the hydrothermally altered rhyolite complex in the Cerro Bayo region, rhyolites on the south side of Cerro Tromen, the wellpreserved biotite-bearing rhyolitic (75%–76% SiO2) ignimbritedome complex and pyroclastic flows at Cerro Tilhué, and the tuffs west of Cerro Tromen (Tilhué Formation on map). The Cerro Bayo, Cerro Tromen, and Cerro Tilhué complexes could be aligned along a northeast- to southwest-trending fault zone. A biotite in a Cerro Bayo rhyolite yielded the new 40Ar/ 39Ar age of 4.04 ± 0.4 Ma, shown in Table 2. Samples analyzed from the Coyocho mafic flows have basaltic andesitic to andesitic compositions and are characterized by arc-like Ba/La ratios (up to 35), La/Ta (24–34), and Ta/Hf (0.17–0.22) ratios, relatively flat REE patterns (La/Yb = 10–16), and moderate Eu negative anomalies (0.62–0.90). Samples from Cerro Tilhué and Cerro Bayo and a pumice clast from a tuff northeast of the Chapúa School are rhyolites with 75%–76% SiO2, 4.3%–4.9% Na2O, and 4.0%–4.7% K2O. Their trace elements show relatively flat REE patterns (La/Yb = 11–16), large negative Eu anomalies (Eu/Eu* is 0.2 in Cerro Tilhué to 0.6 in Cerro Bayo), arc-like Ba/La (18–29) ratios, and intraplate La/Ta (12–17) and Ta/Hf (0.37–0.44) ratios. The pumice is compositionally most like the Cerro Tilhué rhyolite. – Nd (+2.1 to +2.7), and The 87Sr/ 86Sr ratios (0.7041–0.7042), C Pb isotopic ratios in the Cerro Tilhué and Cerro Bayo rhyolites overlap those of younger Tromen region and Southern Volcanic Zone magmas (Figs. 9 and 10). The intermediate-age Chos Malal flows are concentrated in the depression between Cerro Tromen and the Cordillera del Viento that Zapata et al. (1999) designated the Chos Malal trough. Published K/Ar ages for these flows include duplicate ages of 2.3 ± 0.5 and 2.1 ± 0.5 Ma for a flow north of the Chapúa School and an age of 1 ± 0.2 Ma for a flow near Tricao Malal (Valencio et al., 1970). New 40Ar/ 39Ar ages in Table 2 are 1.44 ± 0.08 Ma for the groundmass of a flow near the Chapúa School and 1.04 ± 0.06 Ma for the groundmass of a silicic andesite from Cerro Waile (Fig. 4). As shown in Table 6, most Chos Malal flows have basaltic to basaltic andesite compositions (50%–55% SiO2). Some are characterized by large plagioclase phenocrysts. A sample from Cerro Waile is a silicic andesite. Overall, they are characterized by relatively flat REE patterns (La/Yb = 7–16) with minimal
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heavy REE depletion (Sm/Yb < 3), moderate to small negative Eu anomalies (0.7–0.9), and arc-like La/Ta (27–34), Ba/La (17–24), and Ta/Hf (0.15–0.24) ratios. As shown in Figure 9 and Table 7, Chos Malal mafic flows are slightly less isotopi– Nd near +3.2) than cally enriched (87Sr/ 86Sr = 0.7038–0.7040; C 87 86 – Nd = +2). the Cerro Waile andesite ( Sr/ Sr = 0.7044; C The last group includes the youngest flows from Cerro Tromen and the surrounding region. New analyses in shown in Table 6 and plotted in Figures 5–8 along with those in Llambías et al. (1982) and Stern et al. (1990) show that these flows are primarily andesitic in composition. Duplicate 40Ar/ 39Ar groundmass determinations presented in Table 2 yield an average age of 0.175 ± 0.025 Ma for the young “escorial” flow. Chemically, Cerro Tromen andesites in Table 6 have ~60% SiO2, 2.5%–3% K2O, 1% TiO2, and FeO/MgO ratios ~ 2. Compared to the Cerro Waile Chos Malal trough andesite, they have higher K2O contents, generally less arc-like La/Ta (17–23), Ba/Ta (327–550), and Ta/Hf ratios (0.19–0.2), and higher La/Sm and Sm/Yb ratios. Andesitic blocks in the Agua Carmonina Formation deposits east of Cerro Tromen (Fig. 3) are chemically similar. A basalt (7% MgO, 206 ppm Cr) from south of Cerro Michico (Fig. 3) has distinctly intraplate-like La/Ta (20), Ba/La (19), and Ta/Hf (0.20) ratios and a flat REE pattern – Nd (La/Yb = 6.0, Sm/Yb = 2). The 87Sr/ 86Sr ratio (0.7039), C (+3.4), and Pb isotopic ratios of the “escorial flow” (Table 7; Figs. 9 and 10) overlap those of Southern Volcanic Zone flows to the west. Auca Mahuída and Southern Payún Matrú Fields. Further east are the extensive late Pliocene to Pleistocene Auca Mahuída, Payún Matrú, and Llancanelo alkaline volcanic fields (Fig. 2B). These fields are composed of basaltic to hawaiitic shield volcanoes, monogenetic to polygenetic cones, and differentiated mugearite to trachyandesite flows, domes, and pyroclastic rocks that are concentrated in the high-standing regions. The general characteristics of these fields have been summarized by Bermúdez et al. (1993) and Muñoz Bravo et al. (1989). Only the southern Payún Matrú and Auca Mahuída fields are discussed here. The general distribution of centers in the southern Payún Matrú region is shown on maps by Holmberg (1964), González Díaz (1972, 1979), and Pérez and Condat (1996). The region includes extensive flows in and north of the Sierra de Chachahuén region, as well as the cones and flows near the Río Colorado (Fig. 4). Dated flows have largely yielded Pleistocene ages. Among these are whole-rock K/Ar ages of 1.1 ± 0.5 Ma and 1.0 ± 0.4 Ma for Sierra de Chachahuén flows west of Cerro Ureta (37°2′S, 68°56′W) and north of Cerro Ratón (37°2.11′S, 68°51′W) given in Pérez and Condat (1996). Pérez and Condat (1996) reported a K/Ar age of 2.26 ± 0.07 Ma for a flow from Cerro Tanque in the northern Sierra de Chachahuén. A new 40Ar/ 39Ar groundmass age of 1.48 ± 0.14 Ma for a flow from Cerro Méndez north of the Río Colorado is shown in Table 2. The distribution of centers in the Auca Mahuída field is shown on maps by Holmberg (1964), Ardolino et al. (1996),
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and Delpino and Bermúdez, 2005, personal commun.). The 40Ar/ 39Ar ages in Rossello et al. (2002), which range from 1.7 ± 0.2 Ma to 0.88 ± 0.03 Ma, and new 40Ar/ 39Ar ages in Table 2 show that these rocks erupted in the last 2 m.y., which is in accord with the general Pleistocene age assignments in Holmberg (1964) and Uliana (1978). The groundmass 40Ar/ 39Ar ages in Table 2, correlated with the mapping units of Ardolino et al. (1996), are: 1.78 ± 0.1 Ma for a Pampa de las Yeguas basalt, 1.55 ± 0.07 Ma for a Cerro Las Liebres basalt, 1.39 ± 0.14 Ma for a Cerro Grande basalt, and 0.99 ± 0.04 Ma for a Cerro Auca Mahuída mugearite. The 40Ar/ 39Ar ages are in accord with the relative volcanic sequence proposed by Ardolino et al. (1996) but not the late Miocene to Pleistocene age assignments. The chemistry of southern Payún Matrú and Auca Mahuída field samples analyzed by Kay (2001b) are plotted in Figures 5–9. As in other analyses from this region (e.g., Delpino and Bermúdez, 1985; Bermúdez and Delpino, 1989; Muñoz and Stern, 1988, 1989; Muñoz Bravo et al., 1989; Stern et al., 1990; Bermúdez et al., 1993; Saal et al., 1993, 1995; Saal, 1994), the southern Payún Matrú and Auca Mahuída volcanic rocks are olivine-bearing alkali basalts, hawaiites, benmorites, trachyan-
desites, and trachytes (Fig. 5) characterized by intraplate La/Ta (11–16) and Ta/Hf (0.38–0.52) ratios and transitional arc-like Ba/La ratios (most 16–25). La/Yb ratios range from 8 to 13 in the basalts and from 12 to 18 in more differentiated silicic samples. The higher ratios reflect light REE enrichment in the more silicic flows; all Sm/Yb ratios are <3.5. Overall, the samples show little isotopic variation, with 87Sr/ 86Sr ratios ranging from – Nd from +2.3 to +3.9, and 206Pb/ 204Pb 0.70373 to 0.70388, C from 18.40 to 18.52 (Figs. 9 and 10). DISCUSSION: TECTONIC AND MAGMATIC SYNTHESIS The distribution and chemistry of the magmatic rocks of the Neuquén Basin provide insights on the ages of contractional and extensional deformation across the arc and backarc, the evolution of the underlying crust and mantle, and the geometry of the subducting oceanic plate. The evolution of the Neuquén Basin transect inferred from these data is summarized in cartoon lithospheric cross sections in Figure 11. The cartoons are drawn to scale where possible. Several factors are
Figure 11. (A) Late Cretaceous to Eocene and (B) Miocene to Holocene lithospheric cartoon sections across the Neuquén Basin near 37°S latitude showing model for changing slab dip to explain distribution and sources of magmatism and deformation style. Arrows show contractional and extensional stress regimes. Magmatic centers in white have arc or backarc magmatic geochemical signatures; those in black have intraplate signatures. Percentages are rough estimates of mantle partial-melting percentages. See discussion in text. SVZ—Southern Volcanic Zone.
Upper Cretaceous to Holocene magmatism
Figure 11 (continued).
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important in constructing and interpreting these sections. One is changes in geochemical signals that reflect changes in the mantle and crustal source regions of the magmatic rocks and their implications for the geometry of the underlying continental lithosphere and subducting oceanic plate. Another is constraints from the timing and style of deformation, which are discussed in the following in terms of the magmatic rocks, and relative to other criteria summarized in papers by Ramos and Kay (this volume, chapter 1) and Mosquera and Ramos (this volume, chapter 5). Latest Cretaceous to Paleocene Magmatic and Tectonic History of the Neuquén Basin The characteristics of the Upper Cretaceous to Eocene magmatic rocks in the Neuquén Basin are consistent with their eruption over a relatively steeply dipping subducting plate like that shown in Figure 11. Such a subduction geometry is predicated on Upper Cretaceous to Eocene magmatic rocks between 36°S and 38°S being confined to the western part of the Neuquén Basin (Fig. 2A), the frontal arc-type chemical signatures of the magmatic rocks, and the concentration of concurrent deformation near the frontal arc. The distribution of the magmatic rocks is in accord with the arc front having been east of the Miocene to Holocene arc front. The chemistry of the magmatic rocks is generally similar to that of the Miocene to Holocene arc magmas, as shown by similar high La/Ta, Ba/Ta, and Ba/La ratios and low Ta/Hf ratios (Figs. 6 and 7). Generally, similar La/Yb, La/Sm, and Sm/Yb ratios (Fig. 8) fit with all of these magmas having equilibrated with a similar residual mineral assemblage at a similar temperature and pressure. This commonality suggests that to a first order, the subarc geometry and magma source region were not very different from that today. A lack of backarc magmatism and deformation is consistent with no influence from a subducting slab in the backarc. On the next level, the magmatic and tectonic evolution during this period are not well understood. Available ages indicate long gaps in magmatism, but the distribution and age pattern of the magmatic rocks are only known to a first order. Another factor is that convergence directions between the South American and oceanic plate were oblique for much of this time. Convergence models based on oceanic plates are unconstrained before 69 Ma, show extreme obliquity between 69 and 49 Ma (Pardo Casas and Molnar, 1987), and show lower, but still high degrees of obliquity (>35° in Somoza, 1998) from 49 to 26 Ma. The convergence data permit a more normal convergence regime in the Upper Cretaceous before 69 Ma. The distribution of magmatic rocks is consistent with the arc front having been west of or on the western margin of the Cordillera del Viento from ca. 70 to 50 Ma and shifting to the eastern side at the time of major change in convergence obliquity at ca. 49 Ma. The magmatic record and regional
stratigraphy indicate that the principal period of pre-Miocene contractional deformation in the Neuquén Basin was in the Upper Cretaceous. Evidence for a major Eocene contractional deformation suggested by Cobbold et al. (1999) is less clear. The timing of major periods of deformation is discussed from a magmatic and tectonic viewpoint in the following. Latest Upper Cretaceous (Campanian-Maastrichtian) Support for Upper Cretaceous deformation of Lower Cretaceous and older strata in the western Neuquén Basin and Upper Cretaceous uplift of the Cordillera del Viento comes from magmatic and stratigraphic evidence. Stratigraphic evidence is based on Upper Cretaceous to Paleogene Neuquén and Malargüe Group strata being deposited in a foreland basin east of the Cordillera del Viento (Kozlowski et al., 1987; Barrio, 1990). Magmatic evidence comes from deformed Neuquén Basin strata being cut or overlain by Upper Cretaceous to Eocene magmatic rocks as old as ca. 74 Ma (see Franchini et al., 2003) and by dikes possibly as old as 102 Ma (Zamora Valcarce et al., this volume, chapter 6). A lack of discordance between Upper Cretaceous Neuquén Group and latest Upper Cretaceous– earliest Paleocene Malargüe Group strata to the east shows that deformation was restricted to the general region of the arc in the western Neuquén Basin (Kozlowski et al., 1987). Support for deformation in the latest Upper Cretaceous comes from the biotite 40Ar/ 39Ar cooling age of 69.09 ± 0.13 Ma from the Varvarcó leucogranodiorite in the western Cordillera del Viento (Table 2). Since unrealistically high geothermal gradients are needed to keep biotite from cooling below 300 °C at depths of less than ~3 km (see Kurtz et al., 1997), the present outcrop must have been at 3 km depth or less at 69 Ma. The relatively coarse grain size (up to 0.5 mm) and the amphibole-bearing mineral assemblage of the Varvarcó pluton are consistent with an initial emplacement depth of at least 5–6 km. The case for an Upper Cretaceous emplacement age for the pluton comes from ages of similar magmatic rocks in the Neuquén Basin (see Franchini et al., 2003, and preceding) and from the Lo Valle volcanic rocks to the north in Chile (40Ar/ 39Ar ages of 70–69 Ma; Gana and Wall, 1997). Given an Upper Cretaceous emplacement age for the pluton, 2–3 km of Upper Cretaceous uplift is needed in the Cordillera del Viento region. Contractional deformational at this time fits with a more normal plate convergence regime and evidence for synchronous deformation along much of the Andean margin (see Cobbold and Rossello, 2003). Evidence that the Varvarcó pluton was near the surface by 56.5 ± 0.6 Ma age comes from ages of the Cayanta volcanic rocks (Jordan et al., 2001) that lie unconformably on the western margin of the Cordillera del Viento. A 5–6 km uplift of the pluton between the Upper Cretaceous and Paleocene is consistent with the ~7000 m of pre-Eocene uplift called upon by Kozlowski et al. (1996) to explain the discordance between the
Upper Cretaceous to Holocene magmatism Carboniferous Andacollo Group and the Eocene lavas. Further evidence for Upper Cretaceous deformation comes from fissiontrack ages presented by Burns (2002). Paleocene to Eocene Ages of magmatic rocks in the Neuquén Basin do not provide evidence for the major Eocene contractional deformational event postulated by Cobbold et al. (1999). This event was postulated based on: (1) model ages and orientations of bitumen veins that cut deformed Cretaceous strata on the east side of the Cordillera del Viento, and (2) ages of magmatic rocks in the Sierra de Huantraico (Cobbold and Rossello, 2003). The age for bitumen vein emplacement comes from thermal models that project that organically rich Mesozoic sedimentary source rocks reached thermal maturation in the Eocene. Contemporaneous deformation is predicated on high fluid pressures related to hydrocarbon generation triggering motion on regional detachments. A northwest-southeast orientation for both the veins and major regional contractional structures is considered to reflect tensile failure perpendicular to the least compressive regional stress. This orientation is argued to be the one expected in the right-lateral Eocene transpressional regime predicted from plate convergence vectors (Pardo-Casas and Molnar, 1987). The lack of contemporaneous foreland basin strata is explained by a transpressional regime. Problems for this model are that deformed Neuquén Basin strata are cut by Upper Cretaceous magmatic rocks and that latest Paleocene–earliest Eocene Cayanta lavas unconformably overlie uplifted Cordillera del Viento rocks. The argument for Eocene deformation in the Sierra de Huantraico is complicated by the possibility that the Eocene magmatic rocks used to delimit the deformation age are actually early Miocene in age (Kay and Copeland, this volume, chapter 9). While support for major deformation in the Eocene is unclear, evidence for changes in the magmatic and tectonic regime in the Neuquén Basin as the convergence obliquity changed at ca. 49 Ma comes from the eastward displacement of the arc front. This offset is marked by the cessation of Cayanta volcanism west of the Cordillera del Viento and the initiation of Caicayén group magmatism to the east. The combination of high La/Ta and Sm/Yb ratios in the ca. 50–45 Ma magmatic rocks from Caicayén and Bayo de Esperanza (Figs. 6 and 8) is reminiscent of regionally high ratios in magmatic rocks associated with eastward offsets of the Miocene arc front west of the northern Southern Volcanic Zone (Kay et al., 2005). These Miocene offsets are associated with periods of compressional deformation, and it is reasonable to expect that some Eocene deformation accompanied the eastward shift of the Eocene arc in the Neuquén Basin. Kay et al. (2005) argued that high La/Ta and Sm/Yb ratios at the time of arc migration reflect adjustments in the slab geometry and peaks in forearc subduction erosion (see Fig. 11).
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A Model for Shallowing and Steepening of the Nazca Plate to Explain the Miocene to Holocene Magmatic and Deformational Characteristics of the Neuquén Basin Many aspects of the Miocene to Holocene magmatic and structural evolution of the Neuquén Basin north of the Cortaderas lineament can be explained by shallowing followed by steepening of the subducting Nazca plate as argued by Kay (2001a, 2001b, 2002). Lithospheric-scale cross sections illustrating the model in a transect near 37°S are shown in Figure 11B and are discussed below. They draw heavily from the model for shallowing of the subduction zone below the Chilean (Pampean) flat slab between 28°S and 33°S (e.g., Kay et al., 1991, 1999; Kay and Abbruzzi, 1996).
Early Miocene (Extensional Regime over a Steep Subduction Zone) The lithospheric cross section at ca. 24–20 Ma shows an active volcanic arc and widespread backarc eruptions in an extensional tectonic regime over a relatively steeply subducting slab. The period began as the oceanic Farallón plate broke up and the near-normal subduction regime between South America and the Nazca plate that persists today emerged (see Pardo Casas and Molnar, 1987; Somoza, 1998). Unlike some Oligocene volcanic rocks studied by Muñoz et al. (2000) in the arc region to the south, all of the magmas that erupted in the intra-arc Cura Mallín basin show a frontal arc character as indicated by high La/Ta, Ba/Ta, and Ba/La ratios (Fig. 6). Higher Ta/Hf ratios relative to older and younger arc magmas in the region (Fig. 7) indicate that the mantle wedge had a more intraplate-like character than before or after (Burns, 2002). The largely bimodal basaltic–basaltic andesite and rhyolitic compositional range of the Cura Mallín magmas (Fig. 5) and their relatively flat REE patterns (low La/Yb and Sm/Yb ratios; Fig. 8) are in accord with their eruption through a thin crust in an extensional setting. Further discussion of the setting of the Cura Mallín intra-arc basin can be found in Jordan et al. (2001) and Burns et al. (this volume, chapter 8). In the backarc, a steep subduction zone is consistent with the lack of a subduction-related geochemical signature in the alkali olivine basalts that erupted as far as 550 km east of the modern trench. Their intraplate, OIB-like chemical signatures (see Kay and Copeland, this volume, chapter 9) are shown by low La/Ta, Ba/Ta, and Ba/La (Fig. 6) ratios and high Ta/Hf ratios that reach an extreme in the lavas erupted the furthest east of the arc (Fig. 7). A noncontractional or extensional backarc regime fits with the widespread eruption of alkali olivine basalts. The lack of reported structural evidence for major extension in the backarc appears to suggest that the stress regime was only mildly extensional.
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Late Early Miocene (Change to a Contractional Regime and Initial Shallowing) The lithospheric section at 19–16 Ma shows volcanism from the arc into the middle backarc over a shallower subduction zone than that before 20 Ma. Evidence for a change to a nonextensional stress regime in the arc comes from the end of sedimentation in the Cura Mallín intra-arc basin (Burns et al., this volume, chapter 8). A change from a bimodal volcanic assemblage in the Cura Mallín Formation to a more restricted mafic andesitic composition range in the Trapa Trapa Formation (Figs. 5 and 6) is in accord with a longer residence time for magmas in the crust. A frontal arc setting for the Trapa Trapa lavas is consistent with high La/Ta, Ba/Ta, and Ba/La ratios (Fig. 6). Evidence for a change in the mantle source region comes from a return to typical arc Ta/Hf ratios (Fig. 7). Similarities in La/Yb, La/Sm, and Sm/Yb ratios between Cura Mallín and Trapa Trapa mafic magmas show that no major pressure change occurred in residual mineral assemblages in equilibrium with the erupted magmas. Evidence for a shallower subduction zone under the backarc comes from changes in the character of backarc magmatism. This change is marked by the cessation of widespread alkali olivine basalt eruptions, and their replacement by mafic andesitic to trachydacitic magmas erupting from volcano and caldera complexes in the mid-backarc. Kay and Copeland (this volume, chapter 9) point to the higher La/Ta, Ba/Ta, and Ba/La (Fig. 6) ratios and lower Ta/Hf (Fig. 7) ratios in these midbackarc magmas to argue for introduction of a subducted component into the mantle source below the Sierra de Huantraico by 19 Ma. Evidence that the magmas were more hydrous by 19 Ma comes from the presence of large clinopyroxene phenocrysts in the basalts and hornblende in the mafic andesites. The presence of water is in accord with subduction-related fluids entering the mantle source. The easiest way to explain a more hydrated mantle richer in arc-like subducted components is for a shallower slab to have extended under the Sierra de Huantraico by 20 Ma. The presence of clinopyroxene and amphibole phenocrysts requires a period of magma evolution in the middle to lower crust as would be expected if magma ascent is slowed in a contractional stress regime. Other evidence for a contractional stress regime has already been discussed above. Middle to Late Miocene (Continued Contraction and Shallowing) The lithospheric section at 14–10 Ma shows a shallower subduction zone with volcanic centers in the arc and near backarc and contractional deformation in the backarc. The most important changes are in the backarc. Evidence for little magmatic change in the arc comes from middle to late Miocene (Trapa Trapa, Cajón Negro, and Quebrada Honda) magmas having the same general petrologic character, arc-related La/Ta, Ba/Ta, Ba/La, and Ta/Th ratios (Figs. 6 and 7) and REE ratios (Fig. 8), as early Miocene Trapa Trapa magmas.
Three general stages can be recognized in the middle to late Miocene magmatic history of the backarc. The first is characterized by a virtual magmatic lull from 16 to 14 Ma, the second by andesitic eruptions from ca. 14 to 10 Ma, and the third by another lull from ca. 10.7 to 7 Ma. The backarc magmas of this period have chemical signatures that are consistent with subducted components influencing a mantle wedge above a subducting slab. Among the lavas erupted are the ca. 12 Ma Cerro Negro hornblende andesites, which have high La/Ta (> 30) and Ba/Ta (>1000) ratios and low Ta/Hf ratios that are markedly more arc-like than those of the early Miocene backarc lavas (Figs. 6 and 7). Similarly, the Miocene Huincán magmas in southern Mendoza (Fig. 2A) show clear trace-element evidence for a subducted component (Baldauf, 1997). Support for backarc contractional deformation in the late Miocene comes from chemical signatures in magmatic rocks that are interpreted to reflect relative crustal thickening. The argument is based on REE patterns that have shapes influenced by pressure-sensitive mafic mineral assemblages that equilibrate with mantle-derived magmas in the crust (Hildreth and Moorbath, 1988; Kay et al., 1987, 1991). The best depth indicators are Sm/Yb and La/Yb ratios as they can reflect amphibole that is stable at intermediate pressure and garnet that is stable at higher pressure. Using this reasoning, similar Sm/Yb and La/Yb ratios in ca. 12 Ma Cerro Negro andesites and Eocene Cayanta andesites are consistent with little crustal thickening under the western Neuquén Basin between 56 and 12 Ma. The picture changes in the late Miocene, as shown by La/Yb (Fig. 7A) and Sm/Yb (Fig. 7B) ratios in Miocene to Pliocene Neuquén Basin magmatic rocks. Comparisons show that these ratios are: (1) lower in ca. 14 Ma Huincán I andesites than in ca. 10–6 Ma Huincán II andesites, (2) lower in ca. 12 Ma Cerro Negro andesites than in younger than 4 Ma Tromen andesites, and (3) lower in Cerro Negro andesites than in Huincán andesites. Based on this logic, Baldauf (1997) and Nullo et al. (2002) argued that higher ratios in Huincán II than Huincán I andesites reflected a crustal thickness increase in Mendoza in response to crustal shortening between ca. 14 and 10 Ma. In the same way, lower ratios in ca. 12 Ma Cerro Negro andesites than in Tromen andesites younger than 4 Ma can be interpreted as reflecting crustal thickening in response to crustal shortening between 12 Ma and ca. 4 Ma in Neuquén. Generally higher ratios in Mendoza than Neuquén andesites are consistent with greater crustal thicknesses in Mendoza, where crustal shortening estimates are higher (Zapata et al., 1999). Overall, the data support an episode of backarc crustal thickening between 12 and 10 Ma under the Neuquén Basin. Latest Miocene (Maximum Development of Shallow Subduction Zone) The lithospheric section at ca. 8–5 Ma is drawn as the shallow subduction zone reached its maximum development under the Neuquén Basin. A key feature in this interpretation is the
Upper Cretaceous to Holocene magmatism Chachahuén volcanic complex that erupted in the block-faulted Sierra de Chachahuén, some 500 km east of the modern trench. To the west, the status of ca. 8–5 Ma magmatic activity is poorly known. Candidates for arc magmas at this time are volcanic rocks in the Pichi Neuquén Complex on the west side of the Cordillera del Viento. Analyses of Pichi Neuquén basaltic to basaltic andesites (Table 5; late Miocene in Figs. 5–8) show that they have generally higher La/Ta (>56) and Ba/Ta (>1300) and lower Ta/Hf (<0.9) ratios than the other Miocene arc magmas (compare with Trapa Trapa), as would be expected if they erupted above a dehydrating, shallowly dipping slab. The position of the Pichi Neuquén Complex north of the westward projection of the Cortaderas lineament is consistent with a location west of a shallowly dipping slab. Similarities in REE ratios with other Miocene arc samples (Fig. 8) indicate minimal pressure changes in residual mineral assemblages, as would be expected if crustal thicknesses changed little in the arc region during the Miocene (Fig. 8). The best evidence for a late Miocene shallow subduction zone under the Neuquén Basin comes from the K-rich 7.3– 4.8 Ma Chachahuén volcanic complex in the far backarc (Kay et al., this volume, chapter 10). The high La/Ta and Ba/Ta and low Ta/Hf ratios in the 6.8–4.8 Ma Chachahuén sequences indicate the presence of an arc-like component in their mantle source. This component is most obvious in the late Chachahuén mafic lavas, which have La/Ta and Ba/Ta ratios that approach those of frontal arc lavas (Fig. 6). Evidence for a hydrated, oxidized arc-like magma source also comes from the presence of clinopyroxene, amphibole, Fe-Ti oxides, and titanite phenocrysts. Intraplate-like low La/Ta and high Ta/Hf ratios in ca. 7.3– 6.9 Ma Vizcachas dacites and rhyolites can be explained by the strong influence of an older crustal component on the first erupted silicic magmas (see Kay et al., this volume, chapter 10). Overall, the easiest explanation for the arc-like features in the Chachahuén magmas is introduction of a subducted component into the mantle source over a shallowly dipping slab. This subducted component must be transient, since it is absent in both early Miocene and Pliocene-Quaternary alkali basalts erupted in the Sierra de Chachahuén (Fig. 6). The Cerro Nevado– Plateado centers (Bermúdez, 1991) to the north in Mendoza are considered to have originated in the same way. Further support for a late Miocene shallow subduction zone under the Neuquén Basin comes from similarities in eruption style, setting, and chemistry between the Chachahuén magmas and those in the late Miocene Pocho field, which is located ~700 km east of the modern trench over the Pampean flat slab (Kay and Gordillo, 1994). The shape of the subducting plate in the lithospheric section in Figure 11 is predicated on the Chachahuén magmas erupting ~180–200 km above the slab, as did the Pocho magmas. Mantle melting at this depth in both cases is argued to be related to breakdown of phlogopite and other hydrous phases in the slab (e.g., Poli and Schmidt, 1995). A shallow subduction zone also provides a rationale for a late Miocene episode of contractional deformation that inverts
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older extensional structures leading to uplift of structural blocks across the northern Neuquén Basin. Deformation would be facilitated by hydrous weakening of the underlying lithosphere (Kay and Abbruzzi, 1996), as argued for the Central Andes on geophysical grounds by James and Sacks (1999). Evidence for deformation of the Miocene Trapa Trapa arc sequences and some uplift of the Cordillera del Viento, respectively, comes from structural evidence and fission-track cooling ages discussed by Jordan et al. (2001) and Burns (2002). This deformation must have been over by the middle Pliocene, when the Centinela arc volcanic sequences erupted (Rovere, 1998) west of the Cordillera del Viento. Further east, the Tromen Massif is a high-standing basement block (e.g., Ramos, 1978; Llambías et al., 1982) bounded by relatively steep faults (Zapata et al., 1999). The discussion above is consistent with uplift of the Tromen Massif between ca. 12 Ma and 4 Ma. The most likely times are at ca. 12–10 Ma and during the period of maximum shallow subduction ca. 8–5 Ma. These periods are synchronous with those identified by Baldauf (1997) in Mendoza. The late Miocene uplift of the Sierra de Chachahuén on inverted normal faults is discussed by Kay et al. (this volume, chapter 10). The principal times of Miocene deformation correspond to those of magmatism, as would be expected if hydration is important in triggering both. Pliocene to Holocene (Steepening of the Slab, Extension, and Mantle Melting) The last two lithospheric sections in Figure 11 are for the Pliocene and the Pleistocene to Holocene. Dramatic changes across the Neuquén Basin during this period are interpreted to reflect steepening of the subducting slab north of the Cortaderas lineament. Important changes include mafic volcanism stretching from the arc to the backarc with voluminous eruptions in the far backarc, the end of contractional deformation, and the onset of mild extension. Reasons for tectonic changes in the arc region west and south of the Cordillera del Viento are discussed by Folguera et al. (this volume, chapter 12), Melnick et al. (this volume, chapter 4), and Lara and Folguera (this volume, chapter 14) and references therein. Evidence for extension in that area comes from the formation of the Lancopué trough in a transtensional arc setting (Ramos, 1978; Folguera and Ramos, 2000). A broad NW-SE– trending magmatic belt that developed south of 37°S (e.g., Muñoz and Stern, 1988, 1989) in the late Pliocene narrowed westward into the NE-SW–trending Southern Volcanic Zone by the Pleistocene (Lara and Folguera, this volume, chapter 14). Through these events, the chemistry of Pliocene to Holocene arc magmas west of the Cordillera del Viento retained an arc character with REE patterns similar to older arc magmas in the transect (Figs. 5–8). In contrast, the chemistry of magmatic rocks erupted in the backarc changed dramatically as mafic andesitic to rhyolitic magmas erupted in the Tromen region, and voluminous alkaline
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magmas erupted in the Auca Mahuída and Payenia volcanic fields. In Figure 11, large eruptions in the middle to far backarc are attributed to melting of a hydrated mantle over a steepening subduction zone. Melting of this type would be facilitated by dehydration of the subducting slab, which would produce a mantle “wetspot” during shallow subduction (e.g., Kay and Abbruzzi, 1996; James and Sacks, 1999). Evidence for hydration and incipient melting in the far backarc of a shallow subduction zone comes from a low electrical resistivity anomaly in the mantle under the Chilean (Pampean) flat slab near 32°S (Booker et al., 2004). A return to a steeper subduction angle exposed hydrated mantle with a lowered melting temperature to a hot convecting asthenosphere producing widespread melting. Support for a change to a backarc extensional regime by the Pliocene comes from the Tromen region, where the cross section of Zapata et al. (1999) shows the Chos Malal depression and the Tromen Massif bounded by steeply dipping normal faults. The Zapata et al. (1999) section is consistent with post-Miocene volcanic eruptions in this region being linked to reactivated normal faults. The pre-Pleistocene (Coyocho) flows could be associated with faults along the margins of the Chos Malal trough and the high-Si rhyolites (Tilhué) with faults bounding the Tromen Massif. The rhyolites can be explained by crustal melting in association with injection of mantle-derived basalts. Isotopic similarities between the rhyolites and mafic lavas (Fig. 9) suggest the melted crust was recently underplated by alkaline basaltic magmas. The ca. 4 Ma 40Ar/ 39Ar age of the Cerro Bayo rhyolite (Table 2) shows that extension had begun by 4 Ma. The Pleistocene Chos Malal mafic flows that flowed into the Chos Malal trough between ca. 2 and 1 Ma (Table 2) and covered the Tromen block fit with continued extension. Further east, small eruptions of early Pliocene basalt on the western side of the Auca Mahuída and Payún Matrú fields that preceded the large eruptions in those fields are consistent with initial steepening of the subduction zone in the Pliocene (Fig. 2B). The major eruptions are considered principally Pleistocene in age, based on the 1.8–0.88 Ma span of the 40Ar/ 39Ar ages from the Auca Mahuída field. These ages show that steepening of the subducting slab was well in progress by the Lower Pleistocene. A case for a link between magma eruption and reactivated normal faults can be made from the structural cross section through Auca Mahuída in Zencich (2000). Chemical trends in the Tromen region magmas support a decreasing influence of a subducted component on the backarc magma source as the slab steepened. Evidence comes from Pliocene Coyocho and Pleistocene Chos Malal trough flows being more arc-like in character in that they have higher La/Ta (24–34) and lower Ta/Hf (0.15–0.24) ratios than the Holocene Tromen flows (La/Ta = 17–23; Ta/Hf = 0.19–0.25; Figs. 6 and 7). Furthermore, Pliocene Coyocho flows have higher Ba/Ta (730–1030) and Ba/La (29–39) ratios than younger Chos Malal (Ba/Ta = 430–690; Ba/La = most 16–21; exceptions are the most feldspar-rich flows closest to the arc region) and Tromen (Ba/Ta = 310–510; Ba/La = 17–22) flows consistent with a
larger contribution from fluid mobile slab-derived components in the Pliocene flows (Fig. 6). The Chos Malal flows with the higher Ba/Ta and Ba/La ratios are feldspar-rich flows in the west whose high Ba contents can be attributed to excess feldspar and a location closer to the frontal arc. A weakening arc signature in the Tromen region is consistent with a loss of subducted components in the mantle wedge over a steepening subduction zone. An alternative is that these trends reflect spatial differences in the crust as the Chos Malal lavas are more arc-like than the Tromen andesites. The chemistry of the Tromen andesites would then reflect an intraplate-type crustal contaminant under the Tromen Massif, but not the Chos Malal trough. This contaminant would be like that in the Tilhué rhyolites, or possibly the intraplate-like crustal component in the Vizcachas rhyodacites in the Sierra de Chachahuén (Kay et al., this volume, chapter 10). The argument against this is that 87Sr/ 86Sr isotopic ratios in the Cerro Waile andesite (0.70438) and the Tilhué rhyolites (0.70436–0.70441) are similar, consistent with the same type of crust beneath the entire region. Further support for a changing mantle source over a steepening subduction zone comes from the chemistry of Pliocene to Holocene backarc flows between ~35.5° to 36.5°S in Mendoza (Bermúdez and Delpino, 1989; Bermúdez et al., 1993; Saal et al., 1993, 1995). There, arc signatures are strongest in late Miocene to Pliocene? (Basalt III and IV flows) in the eastern backarc near Cerro Nevado (Ba/Nb ~ 20–65), diminish in Quaternary? (Basalt V) flows in the Llancanelo region (Ba/Nb ~ 20–30), and virtually disappear in the Holocene Los Volcanoes flows (Ba/Nb ~ 15–22) just west of the arc (see Fig. 2B). Saal et al. (1993) interpreted these trends along with increasing incompatible element abundances and decreasing flow volumes as evidence for a decrease in mantle melting in association with the loss of a hydrous arc component in the mantle source, as would be expected over a steepening subduction zone. The strongly intraplate chemistry of the Auca Mahuída and southern Payún Matrú flows, which is indicated by ratios of La/Ta <16 (Fig. 6A) and Ta/Hf up to 0.52 (Fig. 7), is consistent with an intraplate mantle source over a steep subduction zone (Figs. 6 and 7). The greater range of Ba/Ta, Ba/La, and Ta/Hf ratios in these lavas than in the early Miocene intraplate-like Chachahuén and Matancilla basalts (Figs. 6 and 7) can be explained by modification of the mantle by slab-derived fluids during the shallow subduction stage. The range of ratios fits with variable degrees of mixing of fluid and nonfluid modified mantle. The limited range of 87Sr/ 86Sr and εNd values in Auca Mahuída basaltic to trachytic rocks (Fig. 9) is consistent with the more silicic magmas forming by fractionation of the mafic magmas, by partial melting of recently underplated mafic magmas, or both. The greater range in Pb isotopic ratios (Fig. 10) shows some degree of crustal contamination. A relative paucity of silicic eruptions and the minimal evidence for crustal contamination is consistent with a refractory deep crust under the Neuquén Basin. This crust was largely depleted in hydrous,
Upper Cretaceous to Holocene magmatism low-temperature melting components during the extensive melting event that produced the regionally extensive Mesozoic Choiyoi rhyolite complex. The massive amount of late Pliocene to Holocene magmatism in the Neuquén Basin is difficult to explain solely by melting of a mantle “wetspot” over a steepening subduction zone. One explanation is that the Patagonian mantle has remained hot since the Jurassic, and perturbations of an already hot mantle are responsible for large Tertiary to Quaternary plateau basalts in Patagonia (Fig. 1; Kay et al., 2004). The perturbation in this case is a changing subduction geometry. Another factor could be the relative motion of South America over the hotspot reference frame and the effects of rollback. These effects are discussed relative to the origin of the early Miocene magmas in the Neuquén Basin and the large late Oligocene Somuncura plateau province further south (Fig. 1) by Kay and Copeland (this volume, chapter 9). Folguera et al. (this volume, chapter 12) discuss the possible role of slab rollback in explaining the Pliocene to Holocene structural evolution of the arc and forearc in the Neuquén Andes. Significance of the Cortaderas Lineament Another issue is the significance of the northwest-trending Cortaderas lineament, which essentially bounds the southern limit of Neogene backarc volcanism in the Neuquén Basin (Fig. 2). From a Miocene to Holocene viewpoint, the lineament can be interpreted to mark the southern limit of a Miocene shallow subduction zone. Such a model explains: (1) the absence of Neogene backarc magmatism south of the Cortaderas lineament, (2) the changes from an early Miocene non-arc-like chemistry to a late Miocene arc-like chemistry and then back to a Pliocene non-arc chemistry in mafic rocks erupted in the far backarc north of the Cortaderas lineament, and (3) the large volumes of the Auca Mahuída, Payenia, and Llancanelo volcanic fields north of the Cortaderas lineament. On another level, the Cortaderas lineament seems to generally correlate with a change in the structure of the underlying continental crust and mantle lithosphere, since the lineament is near the southern boundary of the Triassic rift systems (e.g., Ramos and Kay, 1991) and marks a general change in Mesozoic to Holocene deformational styles to the north and south (Ramos and Kay, this volume, chapter 1). An open question is the role of changes in the structure of the continental lithosphere in controlling the geometry of the subducting slab. Chemistry of the Arc Front Magmas and the Geometry of the Frontal Arc Lithosphere Another observation is that the trace-element chemistry of the arc region magmas west of and on the eastern margin of the Cordillera del Viento between 36.5°S and 38°S has remained relatively constant from the Paleocene to the Holocene, as would be expected if there was little change in mantle and crustal
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sources over this time. This consistency is shown in the range of the chemical fields in Figures 5 to 8 for magmatic rocks from the Paleocene arc region, the Eocene Cayanta and Caicayén groups, the early Miocene Cura Mallín Formation, the middle to late Miocene Trapa Trapa Formation, the Miocene to Pliocene Cola de Zorro, Cajón Negro, Quebrada Honda, Pichi Neuquén and Centinela Formations, the Quaternary Guañacos Formation and the modern Southern Volcanic Zone arc centers (Planchón, San Pedro, Cerro Azul, Antuco, Llaima, Villarrica and Puyehue). Overall ratios of La/Ta (>25), Ba/Ta (>500), Ba/La (>16), and Ta/Hf (<1.4) are within the range expected in arc rocks. The only values outside of this range are the high Ta/Hf ratios (1.5–1.8) in some early Miocene magmas erupted in the Cura Mallín intra-arc basin. Another key point is that these magmas all have low La/Yb (4–10; Fig. 8A) and Sm/Yb ratios (most <2.5; Fig. 8B), as would be expected in magmas that last equilibrated with low-pressure residual mineral assemblages. The similarity of these ratios is consistent with the crustal and mantle chemistry and geometry beneath the arc having remained relatively constant throughout this time. The similarity of the chemical signatures of the arc magmas near 37°S contrasts with the variations in arc magmas erupted over a similar period to the north. This can be seen near 34.5°S, where the Miocene arc front was displaced to the east between 20 and 16 Ma and again between 8 and 4 Ma (Kay et al., 2005). In that region, La/Yb ratios range from 2 to more than 50, with higher ratios in magmas emplaced during and after eastward shifts of the frontal arc. These arc shifts are inferred to accompany periods of crustal thickening in the arc linked to crustal shortening in the backarc as well as peaks of subduction erosion in the forearc. In comparison, the smaller range of low La/Yb ratios in the arc magmas near 37°S correlates with less crustal shortening in the backarc (e.g., Zapata et al., 1999) and a relatively stable arc front, implying less subduction erosion in the forearc (Fig. 11). The chemistry of the arc rocks fits with fractionation at low pressures and eruption through a thin to normal-thickness crust. The implication is that the arc crust has not been thicker since 56 Ma than the ~40 km that it is today (Yuan et al., this volume, chapter 3). The data are in accord with a model of a relatively constant slab geometry from the trench to the volcanic line, at least at times of frontal arc magmatism, since at least the Paleocene. CONCLUSIONS This overview of the distribution, major- and trace-element chemistry, and isotopic signatures of Upper Cretaceous to Holocene magmatic rocks in an arc to backarc transect across the Neuquén Andes and Neuquén Basin near 37°S has shown that: (1) the character of frontal arc magmatism has changed little since the Upper Cretaceous, and that (2) the early Miocene to Holocene extension of magmatism into the backarc can be reconciled with a transient episode of Miocene shallowing of the subducting Nazca plate. A combination of magmatic and
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structural evidence is consistent with major contractional deformation in the arc in the Upper Cretaceous and contractional deformation across the entire region in the Miocene. A less important contractional event is possible in the early Eocene in association with eastward migration of the arc front. Extensional events are reflected in early Miocene and late Pliocene to Holocene magmas. Overall, the chemistry and spatial distribution of widespread Miocene to Holocene backarc magmatic rocks north of the northwest-trending Cortaderas lineament can be explained by transient shallowing of a segment of the subducting Nazca plate whose southern boundary is marked by the Cortaderas lineament. This shallowing is believed to have begun in the late early Miocene and reached a peak in the late Miocene, coincident with widespread contractional deformation and the eruption of the 7.3 Ma to 4.8 Ma arc-like Chachahuén basaltic to rhyodacitic volcanic complex (also Plateado-Nevado complexes) in the far backarc. The change from an intraplate-like chemistry with no indication of an arc-like component in early Miocene basalts erupted far east of the trench to an arc-like chemistry in late Miocene mafic lavas erupted in the same region attests to the introduction of a slab component in the far backarc magmas. A return to a normal subduction angle and a thicker mantle wedge in the Pliocene would have exposed hydrated mantle to convecting asthenosphere leading to widespread mantle melting. This return is recorded in: (1) the changing arc-like to intraplate-like chemistry of late Miocene to Holocene basaltic to rhyolitic magmas in the Chos Malal trough and Tromen Massif region of the middle backarc, and (2) the extensive late Pliocene to Quaternary intraplate alkaline magmas in the far backarc Auca Mahuída and Payún Matrú volcanic fields. A paucity of silicic eruptions in the backarc is consistent with a refractory crust that lost much of its low-melting component during an extensive early Mesozoic crustal melting event. Chemical similarities in Miocene to Holocene magmatic rocks in the frontal arc region are consistent with (1) a persis-
tent subarc crustal thickness near 40 km, as is the case today (Yuan et al., this volume, chapter 3), (2) a relatively constant mantle wedge geometry, and (3) a similar shape for the subducting plate from the arc front to the trench since the Miocene. Similarities in Paleogene and Miocene to Holocene arc magmas show that such conditions have been present in the trench to arc front region at times of arc magmatism for the last 60 m.y. The lack of changes argues against a major role for forearc subduction erosion in this region since the late Cretaceous. A remaining challenge is to link the Miocene to Quaternary magmatic and tectonic history of the segment of the Neuquén Basin discussed here with the regions to the north and south. If the Cortaderas lineament marks the southern boundary of a transient Miocene shallowly subducting segment of the Nazca plate, does the northern end of the Llancanelo volcanic field near 35.5°S mark the northern boundary? Did subduction of an oceanic plateau cause transient shallowing of the subducting slab (see Kay et al., this volume, chapter 10)? What is the role of changes in lithospheric character in controlling the shape of the subduction zone? ACKNOWLEDGMENTS Funding for this study was provided by a grant from Repsol YPF petroleum company with additional support from U.S. National Science Foundation (NSF) grant 00-87515. The authors thank Repsol YPF for funding and permission to publish the results from: Kay, S.M., Tertiary to Recent Magmatism and Tectonics of the Neuquén Basin between 36.5°S and 38°S Latitude, Final Report to Repsol YPF, April 30, 2001, 215 p. Thanks are also due to D. Ragona, T. Zapata, F. Fuentes, R.W. Kay, and L. Godfrey for variously aiding in the field work, chemical analyses, and processing of satellite images. Discussions with R.W. Kay, V.A. Ramos, T. Zapata, and D. Ragona were helpful in interpreting the data, and reviews by Leopoldo Lopez-Escobar, Charles Stern, and Victor Ramos improved the paper.
Upper Cretaceous to Holocene magmatism APPENDIX 1. AGE SPECTRA FOR 40AR/ 39AR ANALYSES
Figure A1 (on this and following page). Age spectra shown were determined in the laboratory of Peter Copeland at the University of Houston. Analytical procedures can be found in Jordan et al. (2001).
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Figure A1 (continued).
Upper Cretaceous to Holocene magmatism APPENDIX 2. LOCATIONS AND DESCRIPTIONS FOR DATA SAMPLES LISTED IN TABLES 3–6 Samples in Table 3 (Sample Order as in Table) Cretaceous (See Figs. 2A and 3) (1) BPN11. Medium-grained granodiorite from Varvarcó pluton north of town of Varvarcó on the west side of the Cordillera del Viento, 36°49.42′S, 70°40.41′W, 1265 ± 27 m. Cayanta Formation (See Figs. 2A and 4) (2) ESA2. Block from andesitic breccia, aphanitic, dark gray, on road between Andacollo and Guañacos, 37°11.27′S, 70°41.45′ W, 1133 m. (3) ESA7. Dike, aphanitic porphyritic with hornblende phenocrysts, on road between Andacollo and Cayanta, 37°08.34′S, 70°41.93′W, 1310 m. (4) ESA12. Breccia block, aphanitic porphyritic with large amphibole, west of Cayanta on Andacollo-Cayanta road, 37°08.49′S, 70°43.71′W, 1070 m. (5) ESA11. Massive intrusive body, aphanitic, dark gray, on road between Andacollo and Cayanta, 37°08.76′S, 70°39.83′W, 1185 m. (6) ESA5. Possible lava flow, aphanitic with scarce amphibole phenocrysts, on road between Andacollo and Cayanta, 37°08.50′S, 70°42.65′W, 1175 m. (7) ESA3. Dike or sill, aphanitic porphyritic, light gray with asicular hornblende, on road between Andacollo and Guañacos, 37°11.27′S, 70°41.45′W, 1133 m. (8) RG10. Jointed volcanic plug, aphanitic, medium gray, on minor road east of Guañacos in Río Guañacos valley, 37°16.729′S, 70°44.158′W, 1102 m. (9) RG7. Breccia block, aphanitic porphyritic, amphibole and plagioclase phenocrysts, upstream of Guañacos on minor road in Río Guañacos valley, 37°16.228′S, 70°49.102′W, 1287 m. Eocene Cerro Bayo de la Esperanza (See Fig. 2A) (10) TDR8a. Amphibole-bearing basalt in road cut on east side of Route 40 just south of TDR 9, 36°42′S, 69°50.2′W. (11–) TDR9a, b, and c. Cerro Bayo de la Esperanza complex east of Route 40, north of Río Colorado: (a) clinopyroxenebearing basalt flow; (b) hornblende-bearing mafic andesite flow; (c) basaltic clast from volcanic breccia, near 36°41′S, 69°50.45′W. (14) TDR11. Cerro ‘Anteojos’ andesite, road cut south of Barrancas, 36°50.55′S, 69°57′W. (15) DRC16. Cerro de Las Lajas teschenite, from dump outside and south of quarry gate, 36°50′S, 69°3′W. Eocene Cerro Mayal Region (See Figs. 2A and 4) (16) TDR22. Cerro Caicayén andesite near Mina Don Oscar, altered, 37°27.43′S, 70°26.62′W, 1195 ± 30 m.
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(17–19) TDR28a, b, and d. Sierra de Mayal hornblende-bearing mafic andesitic center, samples have greenish cast, all from along road to antenna: (a) at top; (b) near gate, (d) on ledge above gate; near 37°19.91′S, 70°23.69′W, 1807 ± 37 m. (20–23) TDR23a, b, c, and d. Mafic outcrops on both sides of road south of Río Neuquén south of Cerro Mayal: (a) olivine-bearing basaltic andesitic clast from breccia; (b) clinopyroxene-bearing basaltic andesitic clast in breccia; (c) basaltic dike; (d) plagioclase-phyric basaltic andesite dike; near 37°25.15′S, 70°23.18′W, 872 ± 30 m. Miocene Cerro Negro Center (See Figs. 2A and 3) (24) TDR34. Cerro Negro altered subvolcanic andesite, north of Tricao Malal on road to Chapúa, 37°04.69′S, 70°17.91′W, 1606 ± 32 m. (25) TDR21. Cerro Negro andesite flow, west side near puesto (outpost), 37°05.99′S, 70°21.53′W, 1530 ± 25 m. Samples in Table 4 (Sample Order as in Table) Miocene Cura Mallín Formation (See Figs. 2A and 3) (1) VL3. Platy lava flows, aphanitic porphyritic with plagioclase and pyroxene phenocrysts, from outcrop at south end of Vaca Lauquen, west side of lake, 36°51.106′S, 71°04.043′W, 1687 m. (2) RB10. Vesicular andesitic lava flow, north side of Río Buraleo valley, low outcrop east of first high Cura Mallín ridge, 37°00.958′S, 71°00.747′W, 1480 m. (3) LE5. Sill, dark gray aphanitic porphyritic with plagioclase and pyroxene phenocrysts, outcrop near water level at southwest end of Laguna Epulauquen, 36°49.728′S, 71°05.697′W, 1460 m. (4) LE3. Dike, dark gray aphanitic, outcrop on south end of ridge between Vaca Laufquen and Laguna Epulauquen, 36°50.690′S, 71°03.954′W, 1875 m. (5) RB11. Lava flow interbedded with lowermost ridge-forming pyroclastic breccias, north side of Río Buraleo valley, 37°01.472′S, 71°01.394′W, 1462 m. (6) RR6. Graded ash-fall tuff, top of sedimentary Cura Mallín section on south side of Río Reñileuvu, 37°22.27′S, 70°58.44′W, 1443 m. (7) AL2. Granitoid pluton, west side of Arroyo Lumabia, 36°44.304′S, 71°02.603′W, 1838 m. Miocene Trapa Trapa Formation (See Figs. 2A and 3) (8) RLL15. Sill, dark gray aphanitic, immediately downsection from massive porphyritic andesite sill, south side of Río Lileo, 37°13.368′S, 70°55.002′W, 1700 m. (9) RLL13. Andesitic lava flow interbedded with sedimentary rocks, just upsection from columnar jointed sill, south side of Río Lileo above road, 37°12.852′S, 70°55.596′W, 1510 m. (10) RR5. Dike feeding columnar jointed sill/flow, medium gray fine-grained phaneritic, south side of Río Reñileuvu
56
(11)
(12)
(13)
(14)
(15)
(16)
S.M. Kay et al. at base of ridge-forming volcanic section, 37°22.452′S, 70°58.470′W, 1425 m. AP3. Breccia block, dark gray aphanitic porphyritic with plagioclase and pyroxene, west side of Arroyo Palao, north of Río Lileo at base of Cura Mallín–Trapa Trapa transition, 37°08.772′S, 70°59.586′W, 1991 m. RR10. Granitic sill, ridge-forming unit in lower Cura Mallín sequence on south side of Río Reñileuvu, 37°20.544′S, 70°53.478′W, 1115 m. RN4. Breccia block from andesitic flow, near top of slope on east side of road to Lagunas Epulauquen, 36°51.84′S, 70°57.87′W, 1456 m. RLL16. Porphyritic andesite from thick cliff-forming sill on south side of Río Lileo valley, 37°13.368′S, 70°55.002′W, 1700 m. RR9. Dioritic intrusive along Río Reñileuvu road, across river from Moncol pluton, 37°22.068′S, 71°01.152′W, 1545 m. VL4. Granitoid pluton, western end of Vaca Laufquen, 36°53.682′S, 71°06.846′W, 1642 m.
Late Pliocene Cola de Zorro, Centinela Formations (10) RR12. Basaltic flow, columnar jointed, on south side of Río Reñileuvu near junction with road to Andacollo, 37°20.592′S, 70°46.056′W. (11) RR11. Basaltic breccia block from south side of Río Reñileuvu in footwall of main Cura Mallín thrust, 37°20.664′S, 70°53.184′W, 1259 m. (12) RLLC2. Lava flow, black aphanitic with plagioclase and oxidized mafic phenocrysts, Cola de Zorro, Chile, above headwaters of Río Lileo, 37°12.27′S, 71°08.61′W, 1952 m. (13) RLL3. Basaltic lava flow on Argentine side of Buta Mallín pass, above headwaters of Río Lileo, 37°12.88′S, 71°06.99′W, 1961 m. (14) RLL4. Basaltic lava flow along upper portion of Río Lileo, 37°13.17′S, 71°03.97′W, 1622 m. (15) RLL5. Río Lileo, 37°12.91′S, 71°03.19′W, 1688 m. (16) RLL2. Andesitic lava flow above Río Lileo valley north along road toward Andacollo, 37°13.01′S, 70°45.13′W. (17) RLL1. Siliceous pyroclastic flow with pumice fragments along minor road on north side of Río Lileo west from Puesto Arias, 37°13.16′S, 70°45.42′W, 1000 m.
Samples in Table 5 (Sample Order as in Table) Miocene to Pliocene Volcanic Rocks, West of the Cordillera del Viento (See Figs. 2B and 4) (1) BPN11. Cajón Negro Formation andesite, south of town of Varvarcó on east side of river, 36°50.34′S, 70°40.30′W, 1322 ± 43 m. (2) BPN3a. Cajón Negro Formation pumice from ignimbritepyroclastic sequence, west side of road, north of Las Ovejas, 36°50.14′S, 70°41.80′W, 1397 ± 36 m. (3) BPN4. Cajón Negro Formation pumice from ignimbritepyroclastic sequence, road north of Las Ovejas, behind puesto south of Arroyo Ranqui Leo bridge, 36°49.50′S, 70°42.52′W, 1180 ± 40 m. (4) BPN15. Quebrada Honda Formation andesite flow, midway up road to top of plateau on east side of river across from Manzana Amarga, 36°45.62′S, 70°43.62′W, 1567 ± 53 m. (5) SPN1. Pichi Neuquén Complex, dark gray, aphanitic lava flow with mafic phenocrysts in glassy matrix, 30 m below high point on road beside Las Ovejas racetrack, 36°58.146′S, 70°47.298′W, 1561 m. (6) BPN16. Pichi Neuquén Complex clinopyroxene-bearing basaltic andesite flow, north of Pichi Neuquén on road to Laguna de Leche, ~2 km east of gate, 36°33.15′S, 70°45.22′W, 1442 ± 44 m. (7) RN1. Pichi Neuquén Complex glassy dacite, along road on east side of Río Nahueve between Las Ovejas turnoff and Río Buraleo, 37°00.46′S, 70°48.96′W, 1335 m. (8) RN2. Pichi Neuquén Complex platy outcrop, glassy dacite, along road on east side of Río Nahueve just south of Río Buraleo, 36°59.73′S, 70°51.042′W, 1250 m. (9) BPN13. Pichi Neuquén Complex glassy dacite, road to top of plateau on east side of river across from Manzana Amarga, 36°45.68′S, 70°43.40′W, 1673 ± 54 m.
Quaternary to Holocene (18) RG9. Guañacos Formation basaltic andesite, 36°44.304′S, 71°02.603′W, 1838 m. (19) Ant1. Basalt flow from Antuco Volcano along main road in Chile. Samples in Table 6 (Sample Order as in Table) Pliocene Sierra de Huantraico Region (Fig. 2B) (1) HRD5. Northern part of Horqueta Negra basalt flow next to abandoned oil platform, 37°50.3′S, 69°30.7′W. (2) HRD12. Parva Negra basalt flow, 37°41.8′S, 69°34′W. Post-Miocene Cerro Tromen, Tromen Massif, and Chos Malal Trough (See Fig. 4) Older Volcanic Units: Coyocho and Tilhué Formations (3) TDR35. Coyocho basaltic andesite of uncertain age from eroded outcrop under Tilhué tuffs along road south of Chapúa School and north of Tromen turnoff, 37°13.33′S, 70°14.34′W, 1443 ± 40 m. (4) TDR26. Coyocho dark glassy mafic andesite flow in road cut on south side of Route 40, southwest of Cerro Tilhué, 37°19.26′S, 70°9.67′W, 1404 ± 39 m. (5–6) TDR27c and d. Coyocho mafic andesite blocks from breccia at base of Cerro Tilhué, 37°18.77′S, 70°8.41′W, 1709 ± 32 m. (7) TDR24c. Coyocho andesitic breccia clast from fallen block beneath Cero Tilhué rhyolites. 37°18.33′S, 70°7.94′W, 1988 ± 42 m. (8) TDR25. Coyocho andesite flow remnant, south of Cerro Tilhué on road around Estancia Tilhué, near gate to southeast, approximately at 37°22′S, 70°09′W.
Upper Cretaceous to Holocene magmatism
57
(9) TDR17b. Coyocho? andesite from Cerro Bayo region, 37°10.1′S, 69°55.9′W. (10–12) TDR24a, b, and e. Tilhué Formation Cerro Tilhué rhyolite dome complex, samples from large blocks in vega on southwest side in summer grazing area for Estancia Tilhué, 37°18.33′S, 70°7.94′W, 1988 ± 42 m. (a and b) biotite-bearing rhyolite, (e) rhyolite with feldspar and quartz phenocrysts. (13) TDR16. Tilhué Formation Cerro Bayo de Tromen rhyolite, 37°41.8′S, 69°34′W. (14–15) TDR17a and c. Tilhué Formation Cerro Bayo de Tromen rhyolite, 37°10.1′S, 69°55.9′W. (16) TDR20. Tilhué Formation rhyolitic sillar pumice from silicic tuff on east side of road just north of Chapúa school, 37°9.64′S, 70°15.09′W, 1561 ± 35 m.
(29) TDR4. Slightly older Tromen Volcano “escorial” flow (Basalt VI) than TDR2 on northwest side of Cerro Tromen near puesto on west side of Laguna Tromen east of TDR3, 37°06.50′S, 70°05.06′W. (30) TDR2. Tromen Volcano “escorial” mafic andesite flow (Basalt VII) on northeast side of Laguna Tromen, 37°05.50′S, 70°05.5′W.
Basalt III, Chapúa Formation, Chos Malal Trough (See Fig. 4) (17) TDR31. Gray basaltic andesite flow on dirt track west of main road near north end of Los Barros, south of Cerro Waile, 37°5.99′S, 70°8.83′W, 2158 ± 30 m. (18) TDR29. Pinkish massive basaltic andesite flow in outcrop east of road south of Cerro Waile, 37°11.09′S, 70°10.95′W, 1912 ± 56 m. (19) TDR7. Mafic andesite flow just south of Los Barros, 37°08′S, 70°04′W. (20) TDR19. Plagioclase-bearing basalt flow south of Chapúa School on river bank, 37°10.77′S, 70°14.95′W, 1412 ± 62 m. (21) TDR33. Basaltic flow with large plagioclase phenocrysts along road on the westernmost side of Tricao Malal, 37°02.78′S, 70°19.70′W, 1336 ± 30 m.
REFERENCES CITED
Basalt IV, Maipo Formation, Chos Malal Trough (22) TDR1. Basaltic flow near Cerro Mocado east of Cerro Waile and north of Cerro Tromen, 37°03.50′S, 70°01.8′W. Basalt V, El Puente Formation, Chos Malal Trough (See Fig. 4) (23) TDR5. Basaltic andesite from Cerro Waile at end of ski road, 37°03.4′S, 70°08.4′W. (24) TDR30. Dark glassy pyroxene-bearing basaltic andesite flow east of main road on dirt track south of Los Barros to puesto south of Cerro Tromen, 37°10.20′S, 70°9.21′W, 2202 ± 41 m. (25) TDR3. Basaltic andesite flow on northwest side of Cerro Tromen near puesto on west side of Laguna Tromen, 37°06.50′S, 70°05.06′W. (26) TDR6. Silicic andesite on side of Cerro Waile along road to top, 37°03.6′S, 70°09.5′W. Basalt VI and VII, Tromen Formation (See Fig. 4) (27) TDR12. Basaltic Basalt VI flow in valley south of Cerro Michico on north side of road west of Buta Ranquil, 37°03′S, 69°51.60′W. (28) TDR13. Tromen Volcano Basalt VII “escorial” mafic andesite flow south of road just west of Buta Ranquil, 37°03′S, 69°55.73′W.
Agua Carmonina Formation; Age Uncertain (See Fig. 4) (31) TDR15. Mafic andesitic block from agglomerate of uncertain age in Aguada de Carmonina near Casa de Piedra east of Cerro Tromen, south of Buta Ranquil, 37°04.15′S, 69°48.9′W. (32) TDR18. Mafic andesitic block from agglomerate of uncertain age east of Cerro Tromen, 37°09.71′S, 69°57.5′W.
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MANUSCRIPT ACCEPTED BY THE SOCIETY 22 DECEMBER 2005
Printed in the USA
Geological Society of America Special Paper 407 2006
Deep seismic images of the Southern Andes X. Yuan G. Asch GeoForschungsZentrum Potsdam, Telegrafenberg, 14473 Potsdam, Germany K. Bataille Departamento Ciencias de la Tierra, Universidad de Concepción, Concepción, Chile G. Bock* M. Bohm H. Echtler GeoForschungsZentrum Potsdam, Telegrafenberg, 14473 Potsdam, Germany R. Kind O. Oncken GeoForschungsZentrum Potsdam, Telegrafenberg, 14473 Potsdam, Germany, and Freie Universität Berlin, Malteserstraße 74-100, 12249 Berlin, Germany I. Wölbern GeoForschungsZentrum Potsdam, Telegrafenberg, 14473 Potsdam, Germany
ABSTRACT Teleseismic earthquakes recorded within the ISSA (Integrated Seismological Experiment in the Southern Andes) temporary seismic experiment in the Southern Andes between 36° and 40°S latitude have been used to construct receiver function images of the crust and upper mantle. The oceanic Moho of the subducted Nazca plate is observed down to a depth of ~100 km, corresponding well with the WadatiBenioff zone seismicity and wide-angle seismic reflections. Beneath the volcanic arc, the slab begins to be invisible with P-to-S converted waves, implying the completion of the gabbro-eclogite transformation in the oceanic crust at that depth. The continental Moho has been imaged at depth of ~40 km beneath the main cordillera and shallows toward the eastern end of the profile beneath the Neuquén Basin to ~35 km depth. Beneath the Loncopué graben, the Moho is locally uplifted to 30 km depth, possibly resulting from the backarc spreading beginning in the Pliocene-Pleistocene. An anomalously high Poisson’s ratio beneath the volcanic arc may indicate partial melting in the upper-plate crust. Keywords: Southern Andes, subduction zone, P-to-S conversion, receiver function, Moho, slab.
*Deceased (6 November 2002). Yuan, X., Asch, G., Bataille, K., Bock, G., Bohm, M., Echtler, H., Kind, R., Oncken, 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, doi: 10.1130/2006.2407(03). For permission to copy, contact
[email protected]. ©2006 Geological Society of America. All rights reserved.
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INTRODUCTION The Andean orogenic belt, stretching for more than 7000 km along the western margin of South America, was formed during the Cenozoic from convergence of the oceanic Nazca plate and the South American plate. It displays distinct along-strike variations in subduction geometry, arc volcanism, and orogenic shortening (Cahill and Isacks, 1992; Ramos, 1999). Defined by the Wadati-Benioff zone, the dip of the subducted Nazca plate changes several times between normal subduction with angles of 20°–30° and segments of flat subduction with angles of ~10°. The variation of the subduction geometry from north to south is directly related to four active volcanic zones: the Northern (5°N–2°S), Central (16°S–26°S), Southern (34°S–46°S), and Austral (south of 47°S) volcanic zones. All these volcanic zones are located in the normal subduction zones, separated by flat subduction zones without volcanic activity. The Andean orogenic belt varies both in width and elevation among different volcanic zones. In the Central Andean volcanic zone, the orogenic belt is the broadest, with a width of >800 km. Subduction and continental deformation has created two significant plateaus, the Altiplano and Puna at 3.8 and 4.5 km elevation, respectively. The maximum elevation is more than 6000 m. The Wadati-Benioff zone reaches a depth of more than 650 km (Cahill and Isacks, 1992). In the Southern volcanic zone, the mountain belt is much narrower and lower. The mean elevation is <1500 m. The subduction zone seismicity here is much less, and the Wadati-Benioff zone only reaches 200 km depth. These major differences of the surface expression between the Central and the Southern Andean mountain belt point to different architectures of the system and different processes at depth. While the Central Andes have been extensively studied with diverse seismic experiments (e.g., Beck et al., 1996; Yuan et al., 2000; ANCORP Working Group, 2003), little is known about the deep seismic structure of the Southern Andes. Several short industry reflection lines have been imaged mainly the upper crust in the study area (Jordan et al., 2001). Within the collaborative Research Project SFB (Sonderforschungsbereich) 267 “Deformation Processes in the Andes,” an array of broadband seismographs and a network of short-period seismic stations have been deployed in the Southern Andes (Fig. 1). The 450-km-long east-west array roughly along 39°S latitude from the Pacific coast into the Neuquén Basin, consisting of 13 broadband seismic stations equipped with Guralp 3ESP’s and SAM data loggers, was deployed from April 1999 to November 2001. The broadband array was planned for a teleseismic receiver function study. The additional short-period network consisted of 62 seismic stations, equipped with 1 Hz Mark or 5 s MarsLite seismometers and PDAS, MarsLite, and Orion data loggers. It was in operation for ~100 d between January and April 2000, covered an area of 400 × 550 km, and was designed to improve the images of local seismicity and seismic tomography using local earthquakes (Bohm et al., 2002). In addition to the passive-source seismic experiments, a seismic refraction
profile with four shots was executed along the broadband array (Bohm et al., 2002; Lüth et al., 2003). The onshore morphotectonic segmentation of the Southern Andes comprises a forearc, integrating the Coastal Cordillera and the Central Valley, and the Main Cordillera and Neuquén Basin (Fig. 1). The longitudinal depression of the Central Valley contains up to several kilometers of sediments and has been an active extensional depocenter since the late Oligocene (Muñoz et al., 2000; Jordan et al., 2001). The Main Cordillera, with mean altitudes steadily decreasing from ~2700 m at 36°S to <~1000 m at 41°S, is the main morphologic component of the Andean orogen. It includes the southern active volcanic arc (Hildreth and Moorbath, 1988), which is characterized by longlived, stationary (relative to the trench) magmatic activity that has persisted since the Jurassic. The backarc area to the east is dominated by the Neuquén Basin, which formed in a continental embayment as an active depocenter during the Late Triassic to Late Cretaceous (Franzese et al., 2003; Ramos et al., 2004; Ramos and Folguera, 2005;). The basin was inverted during the Late Cretaceous and late Miocene into a fold-and-thrust belt with only moderate shortening compared with the sub-Andean system of the Central Andes (Ramos, 1978). Between the Neuquén Basin and the Main Cordillera, the Loncopué graben formed during the Pliocene to the Quaternary as a result of backarc extension. In this paper, we use the receiver function method to study the shear velocity structure, especially discontinuities, in the crust and upper mantle beneath the seismic stations down to the depth of the mantle transition zone. The broadband station array had a longer operation period and therefore contributed most of the data set. Besides the local earthquakes recorded at the shortterm network stations, a number of teleseismic earthquakes were also recorded within the operation period. The data quality was, however, not sufficient to study the along-strike structural variations. The data processing and interpretation were mainly focused on the broadband array along 39°S latitude. DATA AND METHODS Receiver function analysis has become a routine method for studying crust and upper-mantle discontinuities (Langston, 1977; Vinnik, 1977). It can detect discontinuities beneath seismic stations by identifying P-to-S converted waves at these discontinuities. The converted S waves travel slower to the stations than the direct P wave does, and therefore, will be recorded after the direct P wave in the P wave coda. With the receiver function technique, we can identify these mode conversions and measure the differential time between the converted S and the direct P waves. The weak conversion energy (only a few percent of the P wave energy) can be isolated from the P wave with receiver function analysis, which includes coordinate rotation and deconvolution. Except for the primary conversions (Ps), multiple phases like PpPs and PpSs, which reverberate between the discontinuity and Earth’s surface, can frequently be observed
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Figure 1. Map of stations in the Southern Andes between 36° and 40°S latitude. The recent convergence rate between the Nazca plate and South American plate is 65 mm/yr (Angermann et al., 1999). Slab contours (Gudmundsson and Sambridge, 1998) mark the subducted oceanic lithosphere. Major active faults and structures are included (LOFZ—Liquiñe Ofquí fault zone [FZ]). Young volcanoes are marked by white triangles. The E-W line at latitude 39°S shows the position of the receiver function cross section.
(Fig. 2). The multiples are weaker than the Ps wave because of scattering on their longer ray paths, and so they are sometimes difficult to identify. The multiples have a different distance dependence (moveout) and thus can be distinguished from the Ps phase. Recognition of the crustal multiples helps to accurately estimate crustal thickness and the average crustal Vp/Vs ratio (Zandt et al., 1995; Zhu and Kanamori, 2000; Kind et al., 2002). To calculate receiver functions, teleseismic waveform data of P and PP phases with high signal/noise ratios have been selected. We used P waveforms of earthquakes with epicenter distances ranging from 30° to 95° and magnitudes (mb) larger than 5.5. The PP waveforms have been selected for epicenter distances between 70° and 180° and magnitudes larger than 6.0. Figure 3 shows the location of the 67 earthquakes used in this study. Most of the earthquakes were located in the Central
Figure 2. Ray paths of the directly converted Ps wave (transmission) and the multiple reverberations (reflection) of PpPs and PpSs waves.
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X. Yuan et al. malized to the maximum amplitude of the spike on the L component to preserve the absolute amplitude of conversion ratios. A total of 340 receiver functions were obtained with the broadband array, and a similar amount of receiver functions with the short-period network stations. OBSERVATIONS The Upper-Mantle Discontinuities
Figure 3. Locations of teleseismic earthquakes used in this study. Earthquakes with epicenter distances between 30° and 95° have been selected for P wave analysis. Earthquakes with larger distances (70° to 180°) were used for PP waves.
America and Fiji-Tonga subduction zones, with back azimuths of northwest and southwest, respectively. Receiver functions for each earthquake-station pair were calculated using the method described by Yuan et al. (1997). Three-component seismograms were rotated into an LQT rayoriented coordinate system, in which the L component is in the direction of the incident P wave; the Q component is perpendicular to the L component and is positive away from the source (in SV direction); and the T component is the third component of the LQT right-hand system (in SH direction). The L component is dominated by the P wave, while the Q and T components contain mainly the converted S wave energy. For horizontally layered homogeneous media, the converted S wave energy is exclusively contained in the Q component. The presence of significant energy on the T component indicates laterally varying and/or anisotropic structures. The rotation angles (back azimuth and incidence angle) were determined by the eigenvalues of the covariance matrix over a time window spanning the first few seconds following the P wave arrival. Receiver functions were computed by deconvolving the P waveforms on the L component from the corresponding Q components. A time domain deconvolution procedure was used. First, an inverse filter in the time domain was generated by minimizing the least-square difference between the observed L component and the desired delta-like spike function of normalized amplitude. Then, the inverse filter was convolved with the rotated three-component seismograms. After deconvolution, all components were nor-
The receiver functions have been migrated in an east-west section roughly along the plate motion direction of the subducted Nazca plate (Fig. 4). The amplitudes of receiver functions have been back-projected within one Fresnel zone along the ray path of the incident P wave. The IASP91 model (Kennett and Engdahl, 1991) was used to calculate the time-to-depth transformation. A spatial smoothing window of 20 km was applied over the section to enhance the phase correlation of the upper-mantle discontinuities. Positive amplitudes in Figure 4 (red) mark shear velocity contrast with increasing velocity downward, whereas negative amplitudes (blue) indicate layers with reduced shear wave velocity. Figure 4 shows an overall view of the major features beneath the study area, and the significant events are marked. The continental Moho can be observed by the primary conversions at a depth of ~40 km with lateral topography. A band of energy of multiple reflections within the continental crust can be seen between apparent depths of 120 and 200 km. The subducted oceanic slab, linked with the Wadati-Benioff zone seismicity, can be observed down to a depth of ~100 km. The 660 km upper-mantle discontinuity (the “660”) can be clearly seen at depth of ~650 km. The 410 km discontinuity (the “410”) is weakly visible in the summation trace at a depth of ~400 km. The stability of the observations of the “410” and the “660” is presented in Figure 5, where receiver functions have been stacked in order to compare effects of varied parameters. The signal from 660 km depth can be detected in nearly all of the summation traces. In the majority of cases, converted phases from 410 and 520 km depth can also be seen, but are less pronounced in general. The results indicate that not only are the “410” and “660” stable observations, but the “520” (Shearer, 1990; Helffrich, 2000) is also observed. The 410 and the 660km discontinuities, marking the top and bottom of the mantle transition zone, are generally accepted to be phase changes in mantle mineralogy due to pressure and temperature variations in the mantle. The “410” marks the transformation from olivine to alpha-spinel, and the “660,” the transformation from beta-spinel to perovskite + magnesiowustite (see Helffrich, 2000, for review). Both reactions are sensitive to temperature variation and have Clapeyron slopes of opposite signs. Variations in the discontinuity depths will, therefore, reflect mantle temperature variations; a thicker transition zone is indicative of an anomalously lower temperature, while a thin transition zone would indicate high temperatures. Our data sample the mantle transition zone west of the downgoing slab.
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Figure 4. East-west section of receiver function image across the Southern Andes down to a depth of 800 km. Red color denotes the positive amplitudes (positive velocity jump downward) and blue denotes the negative amplitude (negative velocity jump). The main features have been marked with lines and labels. Solid lines at 410 and 660 km depth mark the global average depths of the 410 and 660 km discontinuities. Circles denote the seismicity between 38° and 40°S latitude. These earthquakes were recorded and located with the short-period mobile network (Bohm et al., 2002). The surface topography along 39°S latitude is plotted at the top. Black triangles mark the broadband stations; red triangles denote the volcanic arc. On the right is shown the summation of the receiver function image, which emphasizes horizontal structures over the entire region.
Because there are no deep earthquakes in this region, teleseismic tomography might help us to define the slab geometry in the deeper part. Although the resolution of the global teleseismic tomography in this region is relatively poor, it is still possible to see that the high-velocity oceanic lithosphere is subducted east of our sampling region at depths of the mantle transition zone (Spakman, 2004, personal commun.). In Figures 4 and 5, both the “410” and the “660” appear to be uplifted by ~10 km (or ~1 s), while the thickness of the mantle transition zone is consistent with the IASP91 model. The apparent discontinuity depth changes should be related to the velocity variations in the upper mantle above the discontinuities. The cold subducted oceanic lithosphere of higher seismic velocity increases the average
upper-mantle velocities, causing the earlier arrivals of the P-to-S conversions of the discontinuities by ~1 s. The thickness of the mantle transition zone is not influenced by velocity heterogeneity in the upper mantle. The normal transition zone thickness indicates that there is no temperature anomaly in the mantle transition zone beneath our array. The Oceanic and the Continental Moho More details can be seen in close-up sections down to a depth of 150 km (Fig. 6). In Figure 4, we observe not only primary converted waves for both the slab and the continental Moho, but also the multiple reflection conversions, as explained
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Figure 5. Summation traces showing converted phases from the mantle transition zone. Theoretical arrival times (in reference to the IASP91 model) of Ps conversions at 410, 520, and 660 km depth are marked by thin lines. (A) Varying frequency filters (bandpass is given on the left in s) have been applied. While the 660 discontinuity broadens, the 410 discontinuity disappears using more longperiod filters. (B) Data set has been reduced by increasing signal-to-noise ratio. A 2–30 s bandpass filter has been applied before stacking. By removing lower-quality data, converted phases become more distinct, in particular the 410 and the 520 discontinuities, which are invisible in the complete data set. (C) Data have been corrected for moveout of the direct converted (Ps) and two multiple (PpPs, PpSs) phases prior to stacking. The upper-mantle discontinuities can be identified in the Pscorrected summation trace. (D) Randomly selected subsets of 267 single traces have been stacked to check the stability of observations (bootstrap method).
in Figure 2. The presence of crustal multiples provides a mean for determining average crustal Vp/Vs along the profiles (Zandt et al., 1995; Zhu and Kanamori, 2000; Kind et al., 2002), further constraining the accuracy of the Moho depth estimate. Similar to Kind et al. (2002), we constructed images using the primary conversions and the multiples. Figure 6A–C shows images of the slab and the continental Moho in which the data has been back-projected assuming all the energy originated
from the indicated phase (Fig. 6A—Ps, Fig. 6B—PpPs, Fig. 6C— PpSs). In each panel, only energy from the indicated phase has been migrated correctly to the approximate places, whereas all other energy would not be correctly migrated. It is clear that both the slab and the continental Moho have been reconstructed by all three kinds of phases (Ps, PpPs, and PpSs), although there are differences in depth among these events. The depth differences among different phases can be attributed to the variations
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Figure 6. East-west section of the crust and the uppermost mantle down to a depth of 150 km migrated by directly converted phases (Ps) (A) and two multiples (PpPs and PpSs) (B and C). Average crustal Vp of 6.0 km/s and Vp/Vs of 1.73 were used. The continental Moho imaged with each wave type is marked in the corresponding panel with different line styles. The solid line on the slab event in each panel marks the direct Ps conversion. In each panel, the energy from the indicated phase should, in principle, have migrated correctly, whereas all other energy should be mispositioned. The observations of the continental Moho from A, B, and C are plotted in D for comparison. If there are no lateral variations in crustal Vp/Vs, the Moho images generated by each phase should mimic each other. A deeper Ps-Moho indicates a higher crustal Vp/Vs ratio, and vice versa.
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in Vp/Vs ratio along the section and the dip of the structures. The back-projection can be biased when dipping interfaces exist. The ray-path deviation for reflected waves caused by dipping interfaces is roughly twice as much as that for transmitted waves; therefore, comparable to the reflection seismics, primary converted waves are less biased by dipping interfaces, whereas errors of migrated multiples at dipping interfaces might be relatively large. This is the reason for the different location of the slab in Figures 6B and 6C. Since the Moho is relatively flat, the difference between main phase and multiples can be interpreted with variations in the Vp/Vs ratio. The subducted Nazca plate can be followed with primary P-to-S converted waves to a depth of 100 km. This phase correlates well with the Wadati-Benioff zone seismicity recorded by the short-period seismic network (Bohm et al., 2002). Coherent slab energy disappears deeper than 100 km, which is also in accordance with the pattern of the earthquake foci, which shows the termination of the intermediate-depth seismicity at the same depth. Similar observations were obtained in the Central Andes (Yuan et al., 2000). We interpret this phase as the oceanic Moho of the downgoing Nazca plate. We picked the times of the continental Moho at the maximum of their conversion signals in Figure 6A–C, and all these “raw” depths are plotted in Figure 6D. The Moho has been continuously imaged by the Ps and PpSs phases, while the PpPs-
Moho only shows two patches with relatively large depth differences compared to the other two phases. In two areas, there are large differences between the Ps- and the PpSs-Moho, where the crustal Vp/Vs ratio is apparently significantly different from the value of 1.73 used in the migrations. Beneath the arc, the Ps-Moho is deeper than the PpSs Moho, while beneath station AS20, located in the Loncopué graben, the Ps-Moho is shallower than the PpSs-Moho. We used a grid-search algorithm introduced by Zhu and Kanamori (2000) to search for crustal thickness and the average Vp/Vs ratio at some sample places. Figure 7 shows results for four subregions: the arc, backarc, the Loncopué graben, and Neuquén Basin. Except beneath the Loncopué graben, where the crustal Vp/Vs ratio is normal to low (1.68), all the other subregions show relatively high Vp/Vs ratios (>1.77). Beneath the arc, the Vp/Vs ratio is extremely high (1.89). The estimated crustal Vp/Vs ratios were used to correct the migrated Moho image (Fig. 8). The results show that the thickness of the Andean crust is ~40 km under the arc cordillera and thins both to the west and to the east. The crust in the Neuquén Basin is approximately 34 km thick. At ~70°W, beneath the Loncopué graben, the crust is thinned by 3–4 km. To the west, the Moho can be traced to ~72.5°W to a depth of 35 km and may be followed to a depth of ~25 km beneath station AS05, where the Lanalhue fault zone is located.
Figure 7. Grid-search for crustal thickness and Vp/Vs ratio by optimally stacking receiver functions in regions of the volcanic arc (A), backarc (B), Loncopué graben (C), and Neuquén Basin (D).
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Figure 8. East-west section of the crust and the uppermost mantle with corrected Vp/Vs ratios along the section shown in the figure. Thick solid lines draw through the maximum amplitudes of the Moho and the slab conversions. The labels on the top mark the tectonic units: CC—Coastal Cordillera, CV—Central Valley, MC—Main Cordillera, and NB—Neuquén Basin. FZ—fault zone.
Crustal Thinning in the Loncopué Trough Beneath station AS20, between 71° and 70°W longitude, the Moho shallows to a depth of ~30 km. The Moho uplift also appears to be visible at station AS19, ~50 km to the west. Grid search analysis of station AS20 shows that the crust is 31 km thick with an average Vp/Vs ratio of 1.68 (corresponding to a Poisson’s ratio of 0.226) (Fig. 7C). In Figure 9A, we show the individual receiver functions of station AS20. The amplitudes at ~2 s (labeled S) relate to a near-surface low-velocity layer (sediments). The converted energy of the Moho is at 3–4 s (M). The PpPs multiple (MM) can be partly seen at a time of 13 s. The energy at 7 s (X) following the Moho conversion may have been caused by intracrustal multiples or an original conversion in the uppermost mantle at 60–70 km depth (an interpretation of similar waveforms was discussed by Mohsen et al., 2004). However, a confident interpretation with the few data presented here cannot be achieved. Figures 9B and 9C show an attempt to model the receiver functions with a simple velocity structure. The model consists of a 31-km-thick crust with a 1.5-km-thick sediment layer. A crustal Vp/Vs ratio of 1.68 was taken from the grid-search algorithm (Fig. 7C). Figure 9C shows a good match between the data (solid line) and the synthetic receiver function (dotted line). The amplitudes at 7 s (X) could be modeled better
when we introduced an additional interface in the crust or the uppermost mantle. However, since we had only one station in the Loncopué graben, this could not be done very reliably. DISCUSSION AND CONCLUSION In Figure 10, we compare the receiver function data with seismic tomography using local earthquakes (Bohm, 2004) and wide-angle reflections (Bohm et al., 2002; Lüth et al., 2003). Receiver functions are normally used to detect interfaces with relatively sharp shear velocity contrast. The mantle lithosphere itself, seen by the low-resolution seismic tomography or surface wave dispersion analysis, is hard to see in receiver functions. When low-seismic-velocity oceanic crust exists on the top of the downgoing lithosphere, the contrast between different materials is increased and thus becomes clearly visible by receiver functions. We observed the oceanic crust to a depth of 100 km. This easterly dipping event correlates well with the wide-angle reflection (Bohm et al., 2002; Lüth et al., 2003), with the WadatiBenioff zone seismicity (Bohm et al., 2002), and with the highvelocity zone in the seismic tomography (Bohm, 2004). Coherent slab energy begins to disappear at a depth of 100 km, as does the intermediate-depth seismicity. At the surface, this location is traced by the active volcanic arc. All this evidence
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Figure 9. Individual receiver functions at station AS20 (A) and a suggested simple model (B) that matches the data well (C). S—Sediment, M—Moho, X—unknown, MM—Moho multiple.
suggests that the dipping structure represents the oceanic Moho of the downgoing Nazca plate. When it is subducted, the oceanic crust is subject to metamorphism and dehydration, which causes earthquakes (Kirby et al., 1996). With increasing pressure, basalt/gabbro in the subducted crust transforms to eclogite. At ~100 km depth, this process is probably completed. Beyond the stability limit of key hydrous phases in the basaltic and gabbroic layers (i.e., amphibole to eclogite transformation), the eclogitization will lead to a comparable seismic velocity as the
surrounding mantle. This causes the oceanic crust to be seismically indistinguishable from the peridotitic mantle. In our data, this transformation appears to be completed at a depth of 100 km, beneath the volcanic arc, which is equivalent to observations in the Central Andes (Yuan et al., 2000). The metamorphism is accomplished by dehydration and fault activation in the oceanic crust, triggering the intermediate-depth earthquakes. Water released from the slab causes partial melting in the overlying mantle and forms the volcanic arc. The extremely high Poisson’s ratio in the crust is consistent with large amounts of fluid and magma accumulation in the crust beneath the volcanoes (ANCORP Working Group, 2003). It is worth noting that a subhorizontal phase with positive polarity (high velocity) at a depth of ~100 km can be clearly seen east of the subducted slab (Fig. 10, red dashed line). The phase can be traced for ~100 km to the east from the place where the slab conversions begin to disappear and the WadatiBenioff seismicity shuts off. Looking at the slab imaged by the multiple reflections in Figures 6B and 6C, a bending of the slab converter at the tip of the slab is remarkable, suggesting that this subhorizontal phase at the end of the slab converter could be a real feature. This interface, however, cannot be the continuation of the slab that bends subhorizontally, because the intermediatedepth seismicity does not match. We speculate that the subhorizontal interface might be a remnant of the old flat slab. From the Late Cretaceous until present, the Nazca plate has changed its subduction angle several times (Ramos and Folguera, 2005; Ramos and Kay, this volume, chapter 1). The last flat subduction history was in the late Miocene (Kay et al., this volume, chapter 2). While the slab was steepening afterward in the Pliocene, a piece of subducted flat slab may have remained within the lithosphere. Alternatively, the interface might represent a hydration front in the mantle wedge, or it may be caused by some hydrous minerals floating away from the slab. The upper-plate continental Moho is easily observed at a depth ranging between 30 and 40 km across much of the section. Compared with the moderate crustal thickness of the Central Andes, the section shows a remarkable Moho topography that can be correlated with the major morphotectonic features at the surface. The crust is ~40-km-thick beneath the arc and the transition to the backarc, and thins to the east in the Neuquén Basin (~35 km), as well as to the west in the transition to the forearc (~35 km). The relative crustal thickening of the Main Cordillera segment (40 km) in relation with the eastern hinterland (34–35 km) fits well with the mean topographic elevation of ~1400 m. Though the Moho farther west is not resolved, crustal thickness values may be as small as 25 km below the Central Valley. However, this transition to the active forearc wedge also coincides with the prominent Lanalhue fault zone (Fig. 1). This inherited pre-Andean and repeatedly reactivated discontinuity correlates with clusters of earthquakes that connect to the slab seismicity (Fig. 6). A series of scattered arrivals can be seen at depths of 25 km, 45 km, and 65 km (Fig. 10), interfering with the slab conversion event. This may indicate
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Figure 10. Comparison of receiver functions with results of seismic tomography and seismic refraction. The seismic tomography at latitude 39°S using local earthquakes (Bohm, 2004) has been plotted in color. Thick line segments denote the wide-angle reflections (Bohm et al., 2002; Lüth et al., 2003). P-wave velocities derived from the seismic refraction data are plotted above and below the wide-angle reflections. Receiver functions have been plotted in grayscale. Only positive amplitudes are shown; red solid lines mark the Moho and the slab by receiver functions; red dashed line marks a mysterious subhorizontal interface. FZ—fault zone.
that the Lanalhue fault zone is a lithospheric-scale feature that cuts through the entire upper plate to the subducted slab. Tomography shows a relatively low P-wave velocity immediately above the slab, which may indicate a serpentinized mantle wedge in the cold forearc domain (Bohm, 2004). However, an unequivocal petrophysical interpretation of the complex boundary zone of the upper and lower plate and the active forearc wedge cannot be made on the basis of the presented receiver function imaging. The wide-angle seismic data was collected in the western part of the profile between 74 and 70°S. The limited data of four shots only show a piece of Moho reflection in the Coastal Cordillera. Elsewhere, the Moho is not seen. An intracrustal interface is detected beneath the Coastal and the Main Cordillera. However, they are not seen in the receiver functions. The absence of the wide-angle reflection of the continental Moho probably reflects the transitional nature of the crust-mantle boundary. A gradient velocity boundary of several kilometers that is detectable by the receiver functions cannot be detected by the short-period wide-angle reflection. Compared to the Central Andes, the average crustal Vp/Vs ratio is apparently higher in the Southern Andes (more than 1.80 in much of the
section). The relatively high Vp/Vs ratio would suggest a large amount of mafic lower crust, which could reduce the velocity contrast between the lower crust and the uppermost mantle. In the eastern part of the section, between 71° and 70°W longitude, the Moho is uplifted by 3–4 km beneath the Loncopué trough in a backarc position. The crustal Poisson’s ratio is 0.226 (Vp/Vs ratio of 1.68), slightly lower than normal. The Loncopué trough is a long depression subparallel to the Cordillera and consists of a half-graben system initiated during the Oligocene and extensionally reactivated in the Pliocene-Pleistocene and active until present (Ramos and Folguera, 2005). Based on this, the observed Moho upwelling is interpreted as an initiating rift structure related to active backarc extension. ACKNOWLEDGMENTS The work was supported by the Sonderforschungsbereich (SFB) 267 of the Deutsche Forschungsgemeinschaft. Seismic instruments were provided by the GeoForschungsZentrum Potsdam (Geophysical Instrument Pool Potsdam), the Freie Universität Berlin, and the Universität Potsdam. We thank Susan Beck and Eric Sandvol for their constructive reviews.
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Geological Society of America Special Papers Neogene tectonic evolution of the Neuquén Andes western flank (37−39°S) D. Melnick, M. Rosenau, A. Folguera and H. Echtler Geological Society of America Special Papers 2006;407;73-95 doi: 10.1130/2006.2407(04)
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Geological Society of America Special Paper 407 2006
Neogene tectonic evolution of the Neuquén Andes western flank (37–39°S) D. Melnick† M. Rosenau GeoForschungsZentrum Potsdam, Telegrafenberg, 14473 Potsdam, Germany A. Folguera Laboratorio de Tectónica Andina, Universidad de Buenos Aires, Buenos Aires 1428, Argentina H. Echtler GeoForschungsZentrum Potsdam, Telegrafenberg, 14473 Potsdam, Germany
ABSTRACT This paper integrates new field observations to summarize the evolution of the 37–39°S segment of the Andean margin during the Neogene period. The western Neuquén Andes represent a transitional segment between the high, broad Central Andes and the low, narrow Patagonian Andes. The Main Cordillera at this latitude was uplifted between 11 and 6 Ma. Since then, extension and transtension has dominated the area. South of 38°S, deformation concentrates along the Liquiñe-Ofqui fault zone, a crustal-scale dextral strike-slip system that accommodates part of the marginparallel component of oblique subduction. The architecture of the volcanic arc is strongly controlled by this fault zone. We differentiate four main tectonic phases: (1) late Oligocene–middle Miocene extension and development of a segmented intra-arc continental rift basin and broad volcanic zone; (2) late Miocene shortening, resulting in uplift, exhumation, and inversion of the former basins and a volcanic gap in the Main Cordillera; (3) Pliocene–early Pleistocene extension of the orogenic structure, reestablishment of the volcanic arc, and transtension along the intra-arc zone; and (4) late Pleistocene–Holocene narrowing of the arc and localized extension-transtension along the axial intra-arc zone. In the Central Andes, shortening has been more or less continuous since the Miocene, whereas in the Neuquén Andes, shortening stopped at ca. 6 Ma, probably related to the increase of the slab angle triggering the extension of the former orogenic structure and the onset of arc-parallel strike-slip faulting. The episodic evolution and migration of volcanism are related to changes in dip of the subducting plate. Keywords: western Neuquén Andes, Main Cordillera, tectonic evolution, arc-parallel fault zones.
†E-mail:
[email protected]
Melnick, D., Rosenau, M., Folguera, A., and Echtler, H., 2006, Neogene tectonic evolution of the Neuquén Andes western flank (37–39°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. 73–95, doi: 10.1130/2006.2407(04). For permission to copy, contact
[email protected]. ©2006 Geological Society of America. All rights reserved.
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INTRODUCTION Along-strike segmentation of convergent margin systems and associated mountain chains has been described worldwide (e.g., Fitch, 1972; Bird, 2003). The Andean convergent margin extends for ~4000 km, with a significant along-strike segmentation of its mountain belt at different scales in time and space (e.g., Gansser, 1973; Jordan et al., 1983; Mpodozis and Ramos, 1989; Dewey and Lamb, 1992; Hervé, 1994; Allmendinger et al., 1997; Jordan et al., 1997; Kley et al., 1999; Lamb and Davis, 2003). Compressional deformation, induced by the convergence between the Nazca and South American plates, is generally accepted as the main driving force that formed the actual Andean orogen, but the timing and onset of deformation remain poorly constrained in some segments of this active margin system. The Neuquén Andes of south-central Chile and Argentina (Fig. 1) record an episodic history of compressional, extensional, and strike-slip tectonic phases and related magmatic activity during the late Cenozoic, contrasting with the Central and Austral Andes, where compression has been more or less steady during this period. This paper integrates new field observations with previous geochronological, stratigraphic, sedimentological, and structural data, enriched by interpretation of digital elevation models
and remote-sensing images, to analyze the timing, styles, and distribution of deformation and its structural control on PlioceneQuaternary volcanism in the western flank of the Chilean Neuquén Andes since the late Miocene. TECTONIC SETTING The Neuquén Andes between 37°S and 39°S represent a transitional domain between the high (>4 km mean elevation) and broad (up to 800 km) Central Andes to the low (mean elevation <1 km) and narrow (~300 km) Patagonian Andes. The Central Andes, north of 34°S reflect differential crustal thickening of up to ~70 km, primarily due to shortening since the Miocene (e.g., Isacks, 1988; Allmendinger et al., 1997; Jordan et al., 1997; Godoy et al., 1999; Giambiagi et al., 2003) and ongoing contractional deformation along the eastern foothills (Cortés et al., 1999). South of 38°S, the Patagonian Andes have a crustal thickness of ~40 km (Bohm et al., 2002; Lüth et al., 2003), and no active foreland fold-andthrust belt has been recognized. Pliocene to Holocene deformation has been localized in the intra-arc zone, which is controlled by the Liquiñe-Ofqui fault zone (e.g., Hervé, 1976; Lavenu and Cembrano, 1999). The subducting Nazca plate beneath the Neuquén Andes consists of ~25- to 35-m.y.-old oceanic crust (Tebbens and
Figure 1. Regional location map, morphotectonic units, and Pliocene-Quaternary faults of the Main Cordillera. Contours indicate the depth to the top of the subducting Nazca plate determined using the local network seismicity from Bohm et al. (2002). Black triangles indicate Holocene volcanoes.
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Neogene tectonic evolution of the Neuquén Andes western flank Cande, 1997). The Nazca–South American plate convergence rate is currently 66 mm/yr, as determined by modeling of global positioning system (GPS) data (Angermann et al., 1999), and ~80 mm/yr averaged over the last 5 m.y. (Somoza, 1998). Major along-strike changes in Andean morphotectonic segments and their tectonic evolution have been attributed to variations in the geometry and physical properties of the downgoing plate (e.g., Jordan et al., 1983; Yañez et al., 2002). In the Neuquén Andes segment, substantial changes in the lower plate occur across the Valdivia fracture zone system, which intersects the margin at ~40°S. Oceanic crust produced by the Chile Rise characterizes the plate to the south, whereas oceanic crust to the north formed at the East Pacific Rise. Crustal age, thickness, number of oceanic fracture zones, and plate rugosity differ markedly across the Valdivia fracture zone (Tebbens and Cande, 1997). The western flank of the Neuquén Andes is subdivided across-strike into five morphotectonic units (Fig. 1): (1) the Coastal Platform, which consists of uplifted Tertiary marine and coastal sequences; (2) the Coastal Ranges, which include a Permian-Triassic accretionary complex and a Paleozoic magmatic arc; (3) the Central Depression, a flat area formed by Oligocene-Miocene sedimentary and volcanic rocks, covered by Pliocene-Quaternary sediments; (4) the Main Cordillera, the focus of this paper, which consists of a long-lived MesozoicCenozoic magmatic arc and intra-arc volcano-sedimentary basins; and (5) the Mesozoic Neuquén Embayment and CretaceousTertiary foreland basin. Liquiñe-Ofqui Fault Zone The Liquiñe-Ofqui fault zone is the dominating structural element of the intra-arc zone of the Patagonian Andes and extends for ~1200 km from 38°S to 46.5°S. This fault zone is a dextral strike-slip system that accommodates about half of the margin-parallel component of oblique subduction (Rosenau, 2004) and deformation resulting from collision of the Chile Rise with the South American continent (Forsythe and Nelson, 1985; Cembrano et al., 2002). South of 44°S, fission-track thermochronology (Thomson, 2002), structural data, and 40Ar/ 39Ar ages of syntectonic micas in mylonitic shear zones (Cembrano et al., 2002) show that the Liquiñe-Ofqui fault zone has acted since ca. 7 Ma as a transpressional zone related to the indention and subduction of the Chile Rise. Between 39.5°S and 42°S, fault kinematic data (Lavenu and Cembrano, 1999; Rosenau, 2004) and 40Ar/ 39Ar ages of syntectonic micas (Cembrano et al., 2000) suggest a dominantly strike-slip regime and both transpressional as well as minor transtensional tectonics within the intra-arc zone since the Pliocene. López-Escobar et al. (1995) noticed that most of the late Quaternary stratovolcanoes and minor eruptive centers south of 38°S are either associated with ~NNE-trending faults of the Liquiñe-Ofqui fault zone or form N50–70°E– and N50–60°W– oriented arc-oblique alignments. Following the model proposed
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by Nakamura (1977), they interpreted the N50–70°E alignments as indicative of the orientation of the axis of maximum horizontal stress (σHmax), and the N50–60°W alignments as pre-existing crustal fractures that also served as channels for magma ascent. South of 38°S, NE orientations of σHmax and strike-slip stress indicators characterize the tectonic regime of the intra-arc zone today (Reinecker et al., 2004). South of ~39.5°S, the Liquiñe-Ofqui fault zone is a prominent fault-zone system expressed morphologically by arc-parallel and arc-oblique fiords and glacial valleys. To the north, in the area of this study, transtensional deformation characterizes the Liquiñe-Ofqui fault zone (Potent and Reuther, 2001; Melnick et al., 2002; Rosenau, 2004). The fault zone loses its morphological expression along the axis of active volcanoes, and its northern termination is formed by several splays described in this study, i.e., the Lonquimay and El Barco fault zones. GEOLOGY OF THE STUDY AREA Based on the available geologic maps (Niemeyer and Muñoz, 1983; Delpino and Deza, 1995; Suárez and Emparán, 1997; SERNAGEOMIN, 2003), the stratigraphy of the western flank of the Neuquén Andes, between 37 and 39°S (Fig. 2), can be divided into seven main sequences: (1) pre-Jurassic volcanic, intrusive, and metamorphic rocks representing the preAndean basement; (2) Jurassic marine and volcanic rift sequences; (3) Mesozoic and Cenozoic intrusive rocks; (4) Upper Cretaceous to Paleogene volcanic and continental sedimentary rocks; (5) late Oligocene to late Miocene volcanic complexes and continental sedimentary rocks; (6) Pliocene to early Pleistocene plateau volcanic and volcaniclastic rocks; and (7) Upper Pleistocene to Holocene stratovolcanoes and minor eruptive centers. Pre-Cenozoic to Paleogene Units The basement of the Neuquén Andes is composed of metasedimentary sequences of late Paleozoic age and PermianTriassic bimodal volcanic and intrusive rocks that form part of a large igneous province known as the Choiyoi Group (Kay et al., 1989). These rocks crop out in the Argentinean Neuquén Andes (Delpino and Deza, 1995) and along the Coastal Ranges (Hervé et al., 1988). Although this pre-Andean basement is not exposed in the Main Cordillera between 37 and 39°S, these rocks are inferred to form the basement as observed immediately to the east in the Cordillera del Viento (70.5°W, 37.2°S) and further north in the Main Cordillera at ~34°S (Giambiagi et al., 2003). Mesozoic and Cenozoic Intrusive Rocks The granitoids that crop out between 37 and 39°S represent the northernmost exposures of the North Patagonian Batholith (e.g., Hervé, 1994; Pankhurst et al., 1999) of Mesozoic to Miocene age. Cretaceous intrusive rocks of the Galletué Plu-
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Figure 2. Generalized stratigraphic chart of the Main Cordillera, divided into northern (37–38°S) and southern (38–39°S) areas. Compiled from (1) Niemeyer and Muñoz (1983), Jordan et al., (2001), Linares et al. (1999); and (2) Suárez and Emparán (1995, 1997).
tonic Group are exposed between 37 and 38°S as an isolated body on the western edge of the Main Cordillera and continuously south of 38.4°S (Fig. 3). This group intrudes the Jurassic units and has a wide age range from 148 ± 8 to 80 ± 2 Ma (K-Ar, biotite, amphibole whole rock; Suárez and Emparán, 1997). Two Paleocene stocks with ages of 63 ± 2 and 58 ± 4 Ma intrude the Cretaceous-Paleogene Vizcacha-Cumilao Complex and are covered by Oligocene-Miocene rocks (Suárez and Emparán, 1997). Gräfe et al. (2002) reported apatite fissiontrack ages of 40.6 ± 4.5 and 33.9 ± 4.6 Ma for these small intrusive bodies. Miocene plutonic rocks are exposed in a N-S–trending, ~5km-wide body between 37 and 38°S. The Melipeuco Plutonic Group (Fig. 3) has ages ranging from 15.1 ± 1.2 to 7.2 ± 1.9 Ma, with an average value of 10.8 ± 1.8 Ma (from the seventeen K-Ar, biotite amphibole ages in Suárez and Emparán, 1997), and intrusion depths are <~3 km (Seifert et al., 2005). Gräfe et al. (2002) reported an apatite fission-track age of 5.8 ± 1.0 Ma for a granite of this group. South of 38.2°S, these rocks generally form the basement of the late Quaternary stratovolcanoes.
Pliocene subvolcanic intrusive rocks are exposed as irregular bodies with ages ranging from 5.2 ± 2.0 to 2.6 ± 0.4 Ma (K-Ar whole rock; Suárez and Emparán, 1997). Oligocene-Miocene Volcanic and Sedimentary Rocks In the Chilean and Argentinean Main Cordillera between 36 and 39°S, sedimentary, volcanic, and volcaniclastic rocks of late Oligocene to middle Miocene age are described in the literature as the Cura Mallín Formation (Niemeyer and Muñoz, 1983; Muñoz and Niemeyer, 1984; Suárez and Emparán, 1995, 1997). Between 37 and 39°S, Carpinelli (2000) and Radic et al. (2002) subdivided the Cura Mallín basin spatially and genetically into two subbasins (Fig. 2), which are described in detail in the respective geographic sections below. These authors, as well as Jordan et al. (2001), concluded on the basis of stratigraphic and structural data and interpretation of seismic reflection lines that the mid-Tertiary basins between 34 and 42°S, including the Cura Mallín basin, were deposited in a continental intra-arc rift during an extensional tectonic regime.
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Neogene tectonic evolution of the Neuquén Andes western flank
Figure 3. Simplified geologic map of the western flank of the Neuquén Andes between 37 and 39°S, compiled from references in the text and our data. Cross sections in Figures 4 and 15 are shown. Areas of Figures 7, 11, 15, and 17 are indicated by boxes. Volcanic centers shown: AN—Antuco–Sierra Velluda, CA—Callaqui, CO—Copahue, AC—Agrio caldera, CM—Cordillera de Mandolegüe, LO—Lonquimay, LL—Llaima, SO—Sollipulli. Faults: LOFZ—Liquiñe-Ofqui fault zone, CAF—Copahue-Antiñir fault, LLFS—Lago de la Laja fault system, RPF—Reigolil-Pirihueico fault.
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Pliocene to Early Pleistocene Plateau Volcanic Rocks Along most of the drainage divide of the Neuquén Andes, Pliocene to early Pleistocene basic volcanic rocks form extensive plateaus. These rocks are referred to as the Cola de Zorro Formation (González and Vergara, 1962; Vergara and Muñoz, 1982) or Asociación Volcánica de la Precordillera Oriental (Suárez and Emparán, 1997) (Fig. 3). The Cola de Zorro Formation overlies the older units in a marked angular unconformity of regional extent. Ages of this unit range from 5.6 ± 0.1 to 1.0 ± 0.1 Ma (K-Ar whole rock; Niemeyer and Muñoz, 1983; Suárez and Emparán, 1997; Linares et al., 1999). Along the western edge of the Main Cordillera, the proximal facies is dominated by lava flows, coarse volcanic breccias, and sills, which grade to the west to a more distal facies characterized by pyroclastic flows and alluvial and fluvial conglomerates assigned to the Malleco Formation (Suárez and Emparán, 1997). This unit forms the current piedmont and upper infill of the Central Depression (Fig. 3), with ages ranging from 4.4 ± 0.5 to 0.8 ± 0.3 Ma (K-Ar whole rock; Suárez and Emparán, 1997). Upper Pleistocene to Holocene Volcanic Centers The late Quaternary Southern volcanic zone extends between 33 and 46°S with an ~NNE trend. The study area is within the central petrological province of this volcanic zone, which is characterized by a constant arc-to-trench distance of ~270 km and products mainly of basaltic to andesitic composition (López-Escobar et al., 1995). In the study region, most of the active volcanic centers occur in linear associations that consist of at least two composite stratovolcanoes and several related alignments of minor eruptive centers. SEGMENTED TECTONO-STRATIGRAPHY AND STRUCTURAL EVOLUTION The stratigraphic descriptions and structural observations of this study are grouped into three geographic areas. These areas represent distinct structural domains with differences in stratigraphy, structural style of deformation, and tectonic evolution. The southern domain is characterized by dominantly strike-slip deformation along the Liquiñe-Ofqui fault zone, the northern domain by extension along the intra-arc and shortening in the Argentinean foothills, and the central domain by a long-lived transfer zone that accommodates the deformation between the northern and southern domains. For the local stratigraphic units of each domain, we follow the terms published by the Chilean Geological Survey. Northern Domain—Lago de la Laja Area (37°S–37.7°S) Stratigraphic Succession Oligocene-Miocene Volcanic and Sedimentary Rocks. The northern subbasin of the Cura Mallín Formation consists of a western part with ~2800 m of late Oligocene to early Miocene
sediments attributed to fluvial, lacustrine, and deltaic-alluvial fan facies, whereas an eastern part records only ~400 m of deltaic sediments (Carpinelli, 2000; Radic et al., 2002). The western succession is characterized by two tectono-sedimentary events, leading these authors to relate this subbasin with a dual cycle of continental rifting. North of 38°S, the Trapa Trapa Formation conformably overlies the Cura Mallín Formation (Niemeyer and Muñoz, 1983). The Trapa Trapa unit is composed of thick pyroclastic breccias, lavas, and minor sedimentary beds. Ages range from 18.2 ± 0.8 to 14.5 ± 1.4 Ma (K-Ar, plagioclase whole rock; Niemeyer and Muñoz, 1983), and in Argentina, an andesitic lava was dated as 16.2 ± 0.2 Ma (Ar-Ar hornblende; Jordan et al., 2001). Upper Pleistocene to Holocene Volcanic Centers. The volcanic centers are: the active Antuco stratovolcano, with a basal lava dated as 0.083 ± 0.04 Ma (K-Ar whole rock; Moreno et al., 1985); a valley-confined basaltic plateau of presumably late Pleistocene age; and the Sierra Velluda volcano, a deeply dissected extinct stratovolcano composed of two units: a lower unit, formed by ~1500 m of lavas, breccias, and intercalated pyroclastic flows, dated as 0.495 ± 0.08 Ma; and an upper unit, formed mainly by ~1000 m of lavas and breccias dated as 0.381 ± 0.04 Ma (K-Ar whole rock; Moreno et al., 1985). The Sierra Velluda and Antuco volcanoes are both emplaced in a circular depression, which might represent a Miocene or early Pliocene caldera or the interference of several basement structures (Fig. 3). Structural Evolution Late Miocene Orogenic Phase: Ñuble Profile. The E-W cross section at ~37.1°S, across the Ñuble National Park, illustrates the bivergent geometry of the orogen in this northern domain (Fig. 4). In the western part, the Huemules thrust (Fig. 5A) is a NNW-striking, moderate- to high-angle east-dipping fault, with ~400 m of slip, which truncates the Cura Mallín and Trapa Trapa Formations. Two synclines flank this thrust. The central part of the Ñuble section is dominated by an open, ~10-kmwide anticline that turns into a syncline at the upper Polcura River (Fig. 4). This syncline is detached by the Calabocillo thrust, which truncates lacustrine strata of the Cura Mallín Formation in the footwall (Fig. 5B). This east-vergent thrust fault, identified by Niemeyer and Muñoz (1983), has a listric geometry and forms a shallow detachment at depth. East of the Calabocillo thrust, the Cura Mallín Formation is deformed in a syncline related to this thrust and an open anticline, which is covered by younger units to the east. At this latitude on the Argentinean foothills, Jordan et al. (2001) presented an industry-style seismic reflection line showing that the Cura Mallín Formation increases in thickness from ~400 to ~1800 m across a west-dipping normal fault. Based on the occurrence of an anticline in the hanging wall of the fault, they interpreted this structure as an inverted, basin-bounding normal fault. Along the western margin of the Main Cordillera, the contact with the Central Depression is an area covered by dense
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Figure 4. Geological cross section along the Ñuble National Park at ~37°S, showing the bivergent geometry of the Andean orogenic structure of the northern domain.
Figure 5. (A) View to the NNW of the west-vergent Huemules thrust. (B) View to the north of the Calabocillo thrust, which has a listric-thrust geometry detaching the Cura Mallín Formation in the hanging wall and truncating the same strata in the footwall.
vegetation and Pliocene-Quaternary sediments and volcanics. This western edge of the Main Cordillera is formed by Cura Mallín strata dipping ~30°W on the flank of a slightly asymmetric west-facing wide anticline. Based on the asymmetry of the anticline and the topographic break, we infer a west-vergent blind thrust underneath this fold. South of the Laja Lake at 37.5°S, a tight, overturned westvergent fold affects sedimentary rocks of the Cura Mallín Formation. The geometry of this fold implies a low-angle thrust
from a detachment in the sedimentary facies. In this northern domain, low-angle structures linked to shallow-level detachments are located in the eastern sedimentary-dominated part of the basin. In contrast, on the western volcaniclastic-dominated part of the basin, steeper faults, such as the Huemules thrust, and wide folds characterize the structure (Fig. 4). Pliocene to Early Pleistocene Extension: Upper Polcura. In the area northeast of Laja Lake (Fig. 3), the NNE-trending Bejar fault (Figs. 6A and 7) juxtaposes Oligocene-Miocene
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Figure 6. (A) View to the southeast of the Bejar fault, juxtaposing Cura Mallín lacustrine strata with Pliocene-Pleistocene volcanic rocks, with at least 300 m of vertical displacement. (B) View to the SSE of the Toro graben. The normal faults offset the unconformity between the Cura Mallín and Cola de Zorro Formations by ~300 m. Location is shown on Figure 7.
lacustrine strata of the Cura Mallín Formation with volcanic rocks from the Cola de Zorro Formation. This fault has at least 350 m of down-to-east dip-slip displacement and extends along strike for ~18 km. West of the Bejar fault, the Toro graben is an elongated, NE-trending depression hanging over the Polcura River (Figs. 6B and 7). The two normal faults offset the angular unconformity between the Cura Mallín and Cola de Zorro Formations by ~300 m vertically. On the western part of the graben, a zone of hydrothermal alteration is recognized along the northern-bounding fault of this structure. The southern fault, on the other hand, is well exposed at a road cut along the upper Polcura River. Late Pleistocene to Holocene Deformation: Laja Region. The Lago de la Laja is a narrow, 32-km-long, volcanic-dammed lake, located along the axis of the intra-arc zone. The Antuco volcano, immediately southwest of the lake (Fig. 7A) suffered a Bandai-san–type caldera collapse event (Lohmar, 2000). The volcanic avalanche emitted during this event and subsequent lava flows dammed the valley, forming the present lake (Vergara and Katsui, 1969). Two Holocene 14C ages have been reported for this event: 9700 ± 600 (Moreno et al., 1985) and 6250 ± 60 yr B.P. (Lohmar, 2000). Seismic-reflection profiles, collected by the RCMG (Renard Centre for Marine Geology, University of Gent, Belgium) (Charlet et al., 2003), show normal faults forming horstand-graben structures cutting the lake-bottom sediments (Fig. 8). These sediments were deposited after the valley was dammed
by the Antuco collapse and are thus younger than 6250 ± 60 yr B.P., indicating recent faulting. The Lago de la Laja fault system (Melnick et al., 2003) runs for ~60 km along the Laja Lake and the Quique and Aguila Rivers (Fig. 7A). The interpretation of a photogrammetric digital elevation model (5 m resolution) and an air photo of the Quique valley (Figs. 7B and 7C) shows a N-S–trending fault scarp that forms an ~2-km-long alignment of dense vegetation due to the concentration of springs along the fractured fault zone. This scarp marks a topographic break at the bottom of the glacial valley (Fig. 7B). Near the junction of the Aguila and Polcura Rivers (Fig. 7A), a wedge-shaped plateau formed by late Pleistocene valleyconfined lavas unconformably covers folded strata of the Cura Mallín Formation (Fig. 7D). A steep topographic break marks the contact between the plateau and the higher Cura Mallín rocks to the east. The plateau morphology, valley confinement, wedge shape, and alignment with the Quique River indicate that the eruption of these flows may have been along the Lago de la Laja fault system. A fault scarp recognized in these lavas indicates that tectonic activity continued after the volcanic eruption (Fig. 7D). Postglacial pyroclastic fallout deposits from the Chillán volcanic complex, located ~30 km to the northwest, cover most of this area. At a road cut south of the plateau, ten N-S–trending normal faults cut these pyroclastic deposits (Fig. 9). The faults form an asymmetric horst-and-graben structure with a westdown polarity. Maximum slip on a west-dipping fault is 3 m. Dixon et al. (1999) reported nine 14C ages of pyroclastic fallout
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Figure 7. Geology of the Lago de la Laja Quaternary fault system. (A) Geologic map with shaded bathymetry from Umweltforschungszentrum Leipzig-Halle, Germany. LLFS—Lago de la Laja fault system. (B) Photogrammetric shaded-relief digital elevation model (5 m resolution); white triangles show the trace of the Lago de la Laja fault system. (C) Air photo of the Quique valley showing alignment of vegetation along the fault. (D) Profile showing the late Pleistocene valley-confined basaltic plateau controlled by the Lago de la Laja fault system.
deposits from the Chillán volcano ranging from 9300 ± 70 to 2270 ± 60 yr B.P. Unfortunately, no datable material was found in the faulted pyroclastics, but we infer that these deposits are equivalent to those dated in the surroundings of the volcano. The horst-and-graben geometry of this outcrop is similar to the structures observed in the seismic line Laja04 (Fig. 8), and their similar age seems well constrained by our observations. On the northwestern slope of the Sierra Velluda, a NWstriking, southwest-dipping normal fault was identified cut-
ting an ~2000 m vertical section of the volcano, with at least 500 m of down-to-the-west vertical displacement. This fault juxtaposes the upper and lower units of the volcano truncating the pyroclastic flows of the lower unit. No caldera collapse events have been described nor identified in the current survey for this volcano. Two NW-striking normal faults affecting Pleistocene lavas were recognized ~20 km west of Sierra Velluda. Therefore, we interpret these faults as late Pleistocene extension.
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Figure 8. The Lago de la Laja fault system in the central part of the lake. Offshore pinger (3.5 kHz) high-resolution seismicreflection profile Laja04, shot by the Renard Centre of Marine Geology, University of Gent, Belgium (Charlet et al., 2003). Line drawing shows normal faults forming horst-and-graben structures. The faults affect sediments that are younger than 6250 ± 60 yr B.P. (14C age from Lohmar, 2000). Location of seismic line is shown on Figure 7. TWT—twoway traveltime; SP—shot point.
Central Transitional Domain— Copahue-Agrio Area (37.7°S–38°S) Stratigraphic Succession Pliocene to Early Pleistocene Volcanic Rocks. In the inner wall of the Agrio caldera, the Cola de Zorro Formation consists of a homogeneous sequence of lavas and breccias deposited between 4.0 ± 0.2 and 5.67 ± 0.2 Ma (K-Ar whole rock; Linares et al., 1999). Ignimbrites and lavas of the Mellizas volcanic
sequence are exposed inside the caldera and surroundings of El Barco Lake in Chile (Melnick et al., 2006), with ages between 2.60 ± 0.2 and 2.68 ± 0.2 Ma (K-Ar whole rock; Linares et al., 1999). The Trolope flows, with ages between 0.8 ± 0.2 and 1.6 ± 0.2 Ma (K-Ar whole rock; Linares et al., 1999), overlie the Mellizas sequence in the northeastern border of the caldera. Pleistocene to Holocene Volcanic Centers. At ~38°S, volcanic activity occurs along the 90-km-long Callaqui-CopahueMandolegüe NE-trending volcanic lineament. The Upper
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Figure 9. Southward view of Holocene fallout pyroclastic deposits affected by normal faults with graben-and-horst geometry. These deposits from the Chillán volcano have 14C ages from 9000 to 2000 yr B.P. (Dixon et al., 1999). Outcrop is located in the upper Polcura River (see Fig. 7).
Pleistocene to Holocene centers emplaced along this structure are: the active Callaqui and Copahue stratovolcanoes, the Agrio caldera, the Trolón minor eruptive center, and the Mandolegüe range. The edifice of the Callaqui stratovolcano defines an 11-km-long, NE-trending morphological ridge (Fig. 10). The Copahue composite stratovolcano, which is located east of the generally NNE-trending volcanic front, has basal lavas with ages of 1.23 ± 0.3 to 0.76 ± 0.2 Ma (K-Ar whole rock; Muñoz and Stern, 1988; Linares et al., 1999) and is marked by numerous eruptions during the Pleistocene and Holocene (Melnick et al., 2006). In the inner part of the Agrio caldera (AC in Fig. 3), isolated subglacial centers of late Pleistocene age are spatially related to WNW-trending extensional faults. In the northern part of the caldera, 1.6 ± 0.2 to 0.8 ± 0.2 Ma lavas and a 0.6 ± 0.2 Ma dome (K-Ar whole rock; Linares et al., 1999) are emplaced along the normal faults, which control the border of the caldera (Folguera and Ramos, 2000). The Trolón eruptive center, located northeast of the Agrio caldera, is dated as 0.6 ± 0.2 Ma at the base (K-Ar whole rock; Linares et al., 1999) and has two Holocene vents. The Mandolegüe range is an elongated fault-bounded block of Pliocene volcanics covered by eroded Pleistocene stratovolcanoes, calderas, and NE-trending dike swarms. Structural Evolution Late Miocene Orogenic Phase: Callaqui-Copahue-Mandolegüe Transfer Zone. Detailed stratigraphic and structural studies (Carpinelli, 2000; Radic et al., 2002) have suggested that the Cura Mallín basin between 37 and 39°S consists of two diachronic depocenters limited by a major transfer zone located at ~37.7°S. Regional surveys focused on PlioceneQuaternary deformation and structural control on volcanism (Folguera and Ramos, 2000; Melnick et al., 2002, 2006) identified an alignment of volcanic activity that extends for ~90 km
Figure 10. Oblique air photo to the northeast, along strike of the Callaqui-Copahue-Mandolegüe lineament. Morphology of the elongated Callaqui volcano can be appreciated. The darker flows are postglacial; on the lower part of the glacier, a fissure can be seen. Note the axial valley on the ridge where small craters and fissures are aligned. The snow-covered hill in the back is the Copahue volcano.
in a N60° direction; it is referred to as the Callaqui-CopahueMandolegüe lineament. The Callaqui-Copahue-Mandolegüe lineament divides two domains with different structural styles, the northern domain, characterized by folding and thrusting of the Oligocene-Miocene units with a bivergent geometry (Fig. 4), and the southern
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domain, where the Mesozoic units are involved in the deformation (more details in the following). Shortening magnitudes, although not quantified in this study, seem to be larger in the northern domain, as inferred from generally tighter folding in this area. Thus, we suggest that the Callaqui-Copahue-Mandolegüe lineament accommodated the difference in total strain and deformation style during the late Miocene shortening phase and that it represents the Pliocene-Quaternary expression of the major transfer zone postulated by Carpinelli (2000) and Radic et al. (2002) for the Cura Mallín basin. Pliocene to Early Pleistocene Transtension: Agrio Caldera. The Agrio caldera is located in the central part of the CallaquiCopahue-Mandolegüe transfer zone at 37.8°S (Fig. 3). This depression has a rectangular morphology and is 20 km long in WNW and 15 km in NE directions. The caldera cuts the Cola de Zorro Formation and the first postcaldera unit is the Mellizas volcanic sequence, which constrains the age of the Agrio caldera to the Upper Pliocene. The Agrio caldera developed between two regional fault zones, the Liquiñe-Ofqui strike-slip and the Copahue-Antiñir thrust systems (LOFZ and CAF in Figure 3, respectively). Around 10 km north of the Agrio caldera, a NE-trending normal fault forms a half-graben with a southeast-down polarity along the Damas River, and ~10 km south of the Agrio caldera along the Chaquilvín River, a north-down normal fault forms another half-graben. These two half-grabens form a symmetrical arrangement with the Agrio caldera in the central part (Fig. 3). The hanging-wall tilt, up to 30°NW in the Damas half-graben, allows the exposure of folded Miocene sequences in the footwall of both faults. Inside the caldera, two WNW-trending graben systems cut the 1.6 to 0.8 Ma Trolope flows, and strike-slip faults cut glacial-polished surfaces and Holocene lavas of the Copahue volcano (Melnick et al., 2006). The rectangular morphology of the Agrio caldera, its spatial association to regional N-S– to NNE-trending faults, and the related structures are compatible with a pull-apart structure that developed during the late Pliocene–early Pleistocene and have been active since then. The Agrio caldera and Callaqui-Copahue-Mandolegüe lineament coincide with the southern limit of the backarc Quaternary shortening observed to the north continuously along the Central Andes (Folguera et al., this volume, chapter 11), and with the northern limit of the strike-slip Liquiñe-Ofqui fault zone that extends along the intra-arc zone from 46.5 to 38°S (e.g., Lavenu and Cembrano, 1999). In the Copahue volcano and adjacent Agrio caldera, Pliocene-Quaternary faulting and the orientation of σHmax, determined from the alignment and morphologies of volcanic effusions (Nakamura, 1977), display heterogeneous patterns reflected in the clockwise rotation of σHmax (Melnick et al., 2006). We interpret this heterogeneous pattern as controlled by the Callaqui-Copahue-Mandolegüe transfer zone, which decouples deformation from the transtensional intra-arc Liquiñe-Ofqui fault zone and the backarc shortening observed along the Copahue-Antiñir fault system north of 37.6°S (Fig. 3) (Melnick et al., 2002). Pliocene-Quaternary volcanism here is concentrated along a crustal-scale, inherited
discontinuity that controls the longest volcanic lineament of the Southern volcanic zone. Late Pleistocene to Holocene Deformation: El Barco. The surroundings of El Barco Lake (Fig. 11) represent the northern end of the Liquiñe-Ofqui fault zone, which forms a series of splays that cut the late Pliocene Mellizas volcanic sequence (Melnick et al., 2006). The splays trend NNW to ENE, diverging from the main NE-oriented Liquiñe-Ofqui fault zone. Glacial incision followed the trend of the main structures, resulting in narrow elongated valleys. Between the Lomín and Chaquilvín Rivers (Fig. 11A), lavas of the Mellizas sequence are cut by a NNE-trending east-down normal fault (Fig. 11C). This fault cuts a glaciated surface, indicating late Pleistocene or younger activity. El Barco Lake is an elongated depression hanging over the Pelahuenco and Treputreo Rivers (Fig. 11B). At a gravel quarry on the southern part of the lake, basaltic lavas overlain by fluvial conglomerates and cross-bedded sandstones dip 40°W; this sequence crops out 300 m higher at the Vizcachas area in a horizontal position. This observation suggests block tilting associated with normal faulting. In the upper Treputreo River, a 3-km-long N-S–trending fault scarp was identified from air photos. The scarp is concave to the east, 35 m high in its central part, and loses topographic relief to the north. This fault is probably the postglacial expression of faults that form the northern termination of the Liquiñe-Ofqui fault zone (Fig. 11A). East of El Barco Lake, fault scarps crosscut postglacial lava flows from the Copahue volcano (Melnick et al., 2006). Southern Domain—Lonquimay Area (38°S–39°S) Stratigraphic Succession Pre-Cenozoic to Paleogene Units. The pre-Andean basement in this area is represented by isolated outcrops at 38.7°S of metamorphic rocks of late Paleozoic–Triassic age, named the Huinucal Ivante strata (Suárez and Emparán, 1997). Unconformably overlying this basement, the Nacientes del Bío Bío Formation (Fig. 3) integrates volcanic and marine sedimentary rocks of Pliensbachian to Oxfordian age (De la Cruz and Suárez, 1997; Suárez and Emparán, 1997), equivalent to the Cuyo Group in Argentina. The Nacientes del Bío Bío Formation is in turn unconformably overlain by the Late Cretaceous–Paleogene Vizcacha-Cumilao complex, which is composed of volcanic and continental sedimentary rocks, including subvolcanic intrusive bodies (Suárez and Emparán, 1997). Oligocene-Miocene Volcanic and Sedimentary Rocks. The southern Lonquimay subbasin of the Cura Mallín Formation is characterized by ~2600 m of early Miocene to middle Miocene sediments, similar to the northern subbasin (~37.5°S) but related to a single tectono-sedimentary cycle of continental rifting (Radic et al., 2002). Suárez and Emparán (1995) determined the main development of the southern subbasin based on 23 ages (K-Ar, biotite whole rock), which are between 22.0 ± 0.9 and 11.9 ± 0.8 Ma. Sedimentological studies show that it was an internally drained closed basin between ca. 18 and ca.
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Figure 11. Late Quaternary normal faulting at El Barco area. (A) Simplified geologic map. Thick lines indicate Pliocene-Quaternary structures, thin lines, late Miocene structures. LOFZ—Liquiñe-Ofqui fault zone. (B) Air photo (left) and structural map over shaded-relief digital-elevation model (right) of El Barco Lake. White triangles indicate postglacial normal faults; gray triangle indicate sag pond. (C) Air photo showing a late Pleistocene and possibly Holocene NNE-trending east-down normal fault.
11 Ma; subsequently, the basin was opened to the regional drainage system, as indicated by deposition of clasts from Cretaceous quartz diorites from the Galletué Plutonic Group and Jurassic graywackes from the Nacientes del Bío Bío Formation (Kemnitz et al., 2005). South of 38.4°S, the Mitrauquén Formation (Suárez and Emparán, 1997) lies conformably above the Cura Mallín unit. The Mitrauquén Formation is formed by thick dacitic ignimbrites, volcanic breccias, andesitic lavas, and interbedded upward-coarsening fluvial conglomerates. Radiometric ages from this unit from Suárez and Emparán (1997) are: four ignimbrites between 8.0 ± 0.5 and 8.3 ± 0.9 Ma, one of 9.5 ± 2.8 Ma at the base (K-Ar biotite), and two andesitic lavas of 8.0 ± 0.3 and 8.1 ± 0.6 Ma (K-Ar whole rock). Upper Pleistocene to Holocene Volcanic Centers. South of 38°S, two main active volcanic centers are related to the NNE-trending Liquiñe-Ofqui fault zone. The first is the Lonquimay volcanic system of Holocene age (LO in Fig. 3), which
is formed by a main cone and a NE-trending, 10-km-long fissure system, including scoria craters, cinder cones, and vents (Moreno and Gardeweg, 1989). North of the Lonquimay volcano, four small minor eruptive centers form another 24-kmlong, NE-trending volcanic alignment. The Llaima volcano is the second volcanic system (LL in Fig. 3), which is composed by a Holocene edifice and a 17-km-long, NE- to NNE-trending field of parasite cones and fissure vents. At 39°S, the Rucapillán volcanic complex is located west of the volcanic front, almost in the Central Depression. This complex is formed by a NNW-trending, 3-km-long alignment of Holocene maars (Moreno and López-Escobar, 1994). To the east, the Sollipulli Pleistocene volcano is composed of a 3-kmwide caldera and several scattered Holocene parasite cones oriented in N-S and E-W directions. This center is located in the eastern part of the volcanic front spatially related to the ReigolilPirihueico fault (Lara et al., 2001), an eastern branch of the Liquiñe-Ofqui fault zone (Fig. 3).
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Structural Evolution Late Miocene Orogenic Phase: Bío Bío Profile. The E-W profile at ~38.8°S, across the upper Bío Bío River (Fig. 12), illustrates the structural style of the southern domain. The Cura Mallín Formation is exposed in a wide anticline and a west-facing flexure. The limits of this uplifted part of the basin are moderate-angle reverse faults that thrust the Mesozoic on top of the Miocene units. In the eastern part of the profile, the N-S–trending Pino Seco thrust (Fig. 13), identified by Suárez and Emparán (1997), superposes the Jurassic Nacientes del Bío Bío Formation on top of the late Miocene Mitrauquén Formation. This structure is cut by a regional angular unconformity at the base of the Pliocene-
Pleistocene Cola de Zorro Formation. Thus, the age of the latest thrusting in this area is between 8.0 ± 0.3 and 4.8 ± 0.5 Ma, which is bracketed by the ages from the Mitrauquén and Cola de Zorro Formations. In the western part of the section, the contact between the Lonquimay Massif and the Cura Mallín Formation is not exposed. This contact was interpreted as an unconformity by Suárez and Emparán (1997) and later as an east-dipping inverted normal fault by Radic et al. (2002). On the eastern side of the upper Pedregoso River (Fig. 12), Cura Mallín strata dip homogeneously ~20°E and shallow eastward, whereas on the western side, folded Mesozoic rocks intruded by Cretaceous and
Figure 12. Geological cross section of the upper Bío Bío River at ~38.5°S. The structure is controlled by the thrusting of the Mesozoic sequence over mildly deformed Oligocene-Miocene Cura Mallín beds.
Figure 13. View to the south of the Pino Seco thrust, juxtaposing Jurassic turbidites of the Nacientes del Bío Bío Formation over ignimbrites and conglomerates from the Mitrauquén Formation. The bedding of the Jurassic strata is subvertical, not clearly seen in the photo. This structure is unconformably overlain by Pliocene-Pleistocene volcanic rocks of the Cola de Zorro F Formation, constraining the last pulse of contractional deformation between 8.0 and 5.6 Ma.
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Neogene tectonic evolution of the Neuquén Andes western flank Miocene granitoids are exposed. Gräfe et al. (2002) reported an apatite fission-track age of 5.8 ± 1.0 Ma for a Miocene granite of the Lonquimay Massif, indicating uplift and exhumation of the massif during the late Miocene. Since the Cura Mallín Formation is only little deformed east of this contact, we interpret this contact as a west-dipping thrust controlling the bulk of the uplift of the Lonquimay Massif (Rosenau et al., 2001). The contact between the Mitrauquén and Cura Mallín Formations is concordant along the eastern Bío Bío Valley. In contrast with the underlying Cura Mallín strata, the dip of the Mitrauquén unit decreases upward from locally 40°E to subhorizontal (Figs. 14A and 14B). The present-day drainage network is formed by torrential rivers, but in contrast, the Mitrauquén conglomerates, which represent braided river systems (Suárez and Emparán, 1997), show evidence of an older landscape with much lower relief. The conglomerate sequence is an ~250-m-thick, coarsening upward unit (Fig. 14B). Along the Mitrauquén Valley, the upper lavas of this unit show a progressive upward-sequence decrease in dip from 20°E at the valley bottom to a horizontal disposition ~300 m higher, where they are conformably overlain by the Cola de Zorro Formation, which suggests syntectonic deposition (Fig. 14C). Thus, we interpret the Mitrauquén Formation as a syntectonic unit, and relate the deposition of the upward-coarsening conglomerate sequence to surface uplift caused by activity of the Pino Seco thrust. A viable scenario is that the Pino Seco thrust was blind during the deposition of the Mitrauquén unit, which was being progressively folded during deposition between ca. 9 and 8 Ma.
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During a later stage, the thrust propagated to the surface, juxtaposing the Nacientes del Bío Bío and Mitrauquén Formations. Erosion and deposition of the Pliocene Cola de Zorro plateau volcanics removed and covered the Mitrauquén unit from the top and part of the back limb of the fold, respectively. Immediately east in the Argentinean foothills, a conglomerate sequence similar to the Mitrauquén Formation unconformably overlies the Mesozoic units. The Zapala basalt, with an age of 8.6 ± 0.4 Ma (K-Ar whole rock; Linares and González, 1990), is intercalated in conglomerates constraining their late Miocene age. These conglomerates may also be related to the uplift of the Main Cordillera at these latitudes. Pliocene to Quaternary Transtension: Lonquimay. The Lonquimay glacial valley (Fig. 15) trends NNE for ~30 km. This linear depression is tectonically controlled and bounded to the west and east by transtensional faults. The following features (Fig. 15) are interpreted from a LandSat ETM+ image and air photos: (1) the Lonquimay River runs along the eastern side of the valley, presumably parallel to a NNE-trending fault system; (2) the fault system is formed by a series of NNE-trending faults, which delimit the occurrence of small ridges in the generally flat valley bottom; (3) triangular and trapezoidal facets characterize the up to 900-m-high western slope of the Lonquimay Massif, which delimits the Lonquimay valley morphologically to the east; and (4) outlets draining the Lonquimay Massif to the west into the Lonquimay valley are dextrally offset by NNE-trending faults. This is observed in fifteen outlets on the eastern part of the Lonquimay valley and on three at the western side of the
Figure 14. Syntectonic deposits of the Mitrauquén Formation. (A) View of steeply dipping fluvial conglomerates. (B) View of subhorizontal, upward-coarsening conglomerates; person on the lower right for scale. (C) View toward the south of andesitic lavas at the upper Mitrauquén valley. Note the upward decrease in dip. Ages are from Suárez and Emparán (1997).
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Figure 15. Quaternary deformation in the Lonquimay transtensional system. (A) Shaded-relief digitalelevation model with main drainage network and Quaternary faults. The Lonquimay valley is interpreted as a hemigraben. Note that the Lonquimay River north of 38.5°S runs on the eastern side of the valley, which may indicate Holocene tilting of the central block. The San Pedro Lake is a sag pond associated with Holocene normal faults shown on Figure 16. (B) Detail of LandSat ETM+ panchromatic band (14.25 m pixel). The white arrows show dextral offset of outlets by NNE-trending faults.
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Neogene tectonic evolution of the Neuquén Andes western flank Punta Negra River. Dextral offset of the outlets along the eastern side of the valley range from ~500 to 200 m. Along a road cut 3.5 km east of Lonquimay, postglacial fallout pyroclastic and alluvial deposits are cut by mesoscale N-S–trending normal faults forming a horst-and-graben structure (Fig. 16). North of this outcrop, a smooth topographic break is observed trending north in the valley floodplain limiting the San Pedro Lake to the east (Fig. 15). We interpret this lake as a sag pond related to the observed Holocene faulting and graben formation in this area. These observations indicate that the Lonquimay area experienced dextral transtensional deformation during the Quaternary, resulting in the formation of a half-graben with east-down polarity and an east-bending pull-apart structure to its northern end (Fig. 15). Dextral offset was accumulated probably during the late Pleistocene–Holocene at this transtensional branch of the northern Liquiñe-Ofqui fault zone. Holocene Volcanism and Deformation: Llaima Cones. The Llaima volcano, located immediately west of the LiquiñeOfqui fault zone (LL in Fig. 3) is one of the largest composite Holocene volcanoes in Chile (Fig. 3). It includes a field of at least 40 parasite scoria cones trending NNE and NE for ~17 km (Naranjo and Moreno, 1991). In the northeastern part of the parasite cone field, interpretation of an air photo shows a set of two en echelon, NE-trending, 400- and 800-m-long faults cutting the side of a postglacial pyroclastic cone (Fig. 17A). These two very recent faults controlled the subsequent effusion of five small vents. However, no clear field observations allowed determination of the kinematics of these faults. According to Tibaldi (1995), in aligned pyroclastic cones that overlie a feeder dike or fault, the breaching angle—the angle between the orientation of the alignment of cones and the
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opening where the lava flows out of the cone (Fig. 17B)—is indicative of the kinematics of the feeder fault in the substratum. In areas under extension, pyroclastic cones usually have faultnormal breaching angles, which are controlled by the dip of the fault and point to the downthrown block. In areas under a transtensional or strike-slip regime, fault-parallel and faultoblique breaching dominates. In the northeastern part of the Llaima volcano, most of the cones have low breaching angles (Fig. 17C), which is indicative of strike-slip or transtensional deformation in the substratum. Following the geological and seismological observations of strike-slip deformation along the Liquiñe-Ofqui fault zone (e.g., Lavenu and Cembrano, 1999; Barrientos and Acevedo, 1992; Rosenau, 2004; Reinecker et al., 2004), we interpret the Holocene NE-oriented alignments of cones and fissures here as extensional shear fractures, which serve as channels for magma. These observations emphasize the control of the Liquiñe-Ofqui fault zone on the volcanic arc. TECTONIC EVOLUTION Based on the integrated structural and stratigraphic data, we establish the following late Cenozoic tectonic evolution of the western Neuquén Andes between 37 and 39°S (Fig. 18). Cura Mallín Basin Formation: Late Oligocene–Middle Miocene (ca. 28–11 Ma) During this stage, extensional tectonics and basin formation was contemporaneous with an increase in convergence rate between the Nazca and South American plates (Muñoz et al., 2000; Jordan et al., 2001). The Cura Mallín volcano-sedimentary
Figure 16. View to the south of Holocene fallout pyroclastic and alluvial deposits affected by normal faults with graben-and-horst geometry. Outcrop along international road, 3.5 km east of Lonquimay (see Fig. 15).
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Figure 17. Holocene volcanism and tectonics at the Llaima volcano. (A) Orthorectified air photo of the northeastern part of the Llaima volcano. (B) Geologic and topographic map of the same area as A. Contour interval is 50 m. (C) Histogram of the breaching angles measured in B (the breaching angle is the angle between the orientation of the alignment of cones and the opening from where the lava flows out of the cone). According to Tibaldi (1995), low breaching angles are indicative of strike-slip deformation along a feeder fault in the substratum. (D) View to the east of three aligned pyroclastic cones. The arrows indicate the direction of breaching. The Liquiñe-Ofqui fault zone (LOFZ) and Melipeuco granite (late Miocene) are shown. Note the pronounced scarp associated with the Liquiñe-Ofqui fault zone.
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Neogene tectonic evolution of the Neuquén Andes western flank intra-arc basin was formed (Burns et al., this volume, chapter 8) as a continental rift with lacustrine sedimentation in internally drained subbasins with volcaniclastic input from the west (Suárez and Emparán, 1995; Kemnitz et al., 2005). The Cura Mallín basin between 37 and 39°S is divided along-strike into two diachronic subbasins limited by the NE-trending CallaquiCopahue-Mandolegüe transfer zone (Carpinelli, 2000; Radic et al., 2002; Melnick et al., 2002). Shortening and Uplift: Late Miocene (11–6 Ma) Following extension, a compressional tectonic regime during the late Miocene occurred coeval with the emplacement of large amounts of granitoid melts into the crust. During this compressional stage, the present-day Main Cordillera was built. Shortening resulted in uplift and exhumation of granitoids that intruded at depths of less than 3 km (Seifert et al., 2005). Deformation in the northern structural domain (37–37.7°S) was characterized by a bivergent wedge geometry, similar to the mid-Tertiary basin at ~34°S (Godoy et al., 1999), although shortening amounts were smaller. Thrusts with a shallow detachment level developed in the eastern, sedimentary part of the basin, while the western volcaniclastic part was characterized by high-angle reverse faults. Deformation style differed in the southern structural domain (38–39°S), where uplift of Mesozoic units occurred along reverse faults. Shortening magnitudes decrease from north to south in the study area. The Callaqui-Copahue-Mandolegüe transfer zone marks the boundary between the northern and the southern structural domains accommodating the differences in total strain. The onset of mountain building, relief formation, and exhumation at ca. 11 Ma in the southern subbasin is constrained by clast-provenance studies (Kemnitz et al., 2005). The 9.5 to 8.0 Ma conglomerates and lavas of the Mitrauquén unit are interpreted as syntectonic deposits of this orogenic phase. Ignimbrites and minor lavas of the Mitrauquén unit are the only volcanic products of late Miocene age in the Main Cordillera at this latitude, where shortening ceased before 5.6 Ma. This volcanotectonic scenario is similar to the one proposed by Coira et al. (1993) for the ignimbrites to the north in the Puna region. Liquiñe-Ofqui Fault Zone Transtension: Pliocene–Early Pleistocene (5–1 Ma) Cessation of shortening in the area was accompanied by migration of the volcanic arc from an eastern position at ~69–70°W toward the trench over more than ~150 km (Kay et al., this volume, chapter 2). Volcanism was reestablished in the Main Cordillera, in an ~50-km-wide zone north of 38°S, which may be up to ~90 km wide to the south. Reestablishment of the volcanic arc was temporally coincident with the renewal of extensional deformation. During this stage, as the obliquity of plate convergence was continuously increasing (Somoza, 1998), motion on the Liquiñe-Ofqui fault zone initiated. Transtensional
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deformation was localized along the northern segment of the Liquiñe-Ofqui fault zone, including the Agrio caldera pull-apart structure and the Lonquimay and El Barco fault zones. Continued slip along the Callaqui-Copahue-Mandolegüe structure controlled the strain transfer from the transtensive northern Liquiñe-Ofqui fault zone to the Copahue-Antiñir thrust front in the backarc flank of the Main Cordillera. Extension and Transtension: Late Pleistocene–Holocene (1–0 Ma) During the late Pleistocene, the volcanic arc narrowed to its present position (Stern, 1989). Deformation is closely associated with volcanic activity along the Liquiñe-Ofqui fault zone in the axial intra-arc zone. All the centers south of the CallaquiCopahue-Mandolegüe transfer zone have Liquiñe-Ofqui fault zone–related strike-slip structural elements. North of this transfer zone, late Pleistocene–Holocene deformation is localized along the Lago de la Laja extensional fault zone and to the north of ~37.5°S compressional deformation is continuously present along the backarc (Folguera et al., this volume, chapter 11). DISCUSSION: GEODYNAMIC IMPLICATIONS The factors controlling extension and shortening in the Southern Andes have been the subject of controversial debate. The late Oligocene plate reorganization in the South Pacific (Tebbens and Cande, 1997) produced an increase in the convergence rate between the Nazca and South American plates (Somoza, 1998). Muñoz et al. (2000) and Jordan et al. (2001) showed that conventional models linking shortening in the overriding plate to high convergence velocities do not account for the Neogene tectonic evolution of the Neuquén Andes, since regional extensional tectonics and basin formation are coeval to high convergence rates in the Andes south of ~33°S. The onset of shortening in the Neuquén Andes during the late Miocene is coincident with the flattening of the slab, emplacement of shallow crustal granitoids, and a lack of volcanism in the Main Cordillera, as discussed by Kay (2002) and Kay et al. (this volume, chapter 2). Low subduction angles and the onset of shortening in the Miocene have been previously proposed for the Central and Southern Andes between ~19 and 29°S (Isacks, 1988; Allmendinger et al., 1997). Shortening in the Neuquén Andes ceased at ca. 6 Ma and was followed by the reestablishment of volcanic activity in the Main Cordillera and extensional tectonics related to steepening of the slab and a westward trench retreat (Muñoz and Stern, 1988; Stern, 1989). This scenario contrasts with those proposed for the areas north of ~34°S, where the downgoing plate maintained a lower angle and shortening has been more or less continuous since the Miocene and is still ongoing in the foreland (Giambiagi et al., 2003; Cortés et al., 1999). The steepening of the slab in the Neuquén Andes is considered to be the main driving factor for the cessation of shortening and the onset of extension during the Pliocene.
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Neogene tectonic evolution of the Neuquén Andes western flank Two factors may explain the onset of the strike-slip LiquiñeOfqui fault zone at ca. 6–5 Ma: (1) the plate reorganization, which caused a continuous increase in the obliquity of the Nazca–South America convergence since the late Miocene (Somoza, 1998); and (2) steepening of the slab and reestablishment of the volcanic front along the Main Cordillera, which causes thermal weakening and strain localization in the intra-arc zone. Arrival of the northward migrating Chile Rise collision zone at the southern end of the volcanic arc might have additionally triggered initiation of the Liquiñe-Ofqui fault zone (Forsythe and Nelson, 1985). Localization and evolution of the northern end of the Liquiñe-Ofqui fault zone at 38°S may be related to a combination of the following factors: (1) along-strike changes in physical properties of the oceanic plate at the Valdivia fracture zone system, separating a segmented and rough oceanic crust produced by the Chile Rise from a homogeneous and smoother crust produced by the East Pacific Rise (Tebbens and Cande, 1997); (2) widening of the orogen northward by increasing backarc shortening during the Miocene (Kley et al., 1999) and partitioning of oblique subduction over a broad area of deformation in the Central Andes (Folguera et al., 2002); and (3) a probable threshold distance in the ridge-push force exerted by the Chile Rise northward (Cembrano et al., 2000). CONCLUSIONS The western flank of the Neuquén Andes between 37 and 39°S records an episodic evolution during the Neogene, which can be summarized in four main tectonic phases: (1) late Oligocene– middle Miocene extension and development of a wide zone of volcanic activity and a segmented intra-arc continental rift basin; (2) late Miocene shallowing of the slab, triggering a gap
Figure 18. Series of interpretative maps summarizing four main phases in the tectonic evolution of the western Neuquén Andes between 37 and 39°S, based on structural data, stratigraphic facies, sediment thickness, interpretation of digital-elevation models and remote-sensing data, and published information. The town of Lonquimay is represented by a black square. (A) Formation of the Cura Mallín basin during extensional tectonics, CCM—Callaqui-Copahue-Mandolegüe transfer zone; long stippled lines indicate inferred secondary transfer zones. (B) Inversion of the Cura Mallín basin, exhumation of Miocene granitoids, and hiatus in volcanism except for the Mitrauquén ignimbrites. (C) Reestablishment of the volcanic arc, extension of the former orogenic structure, transtensional deformation associated with the northern Liquiñe-Ofqui fault zone (LOFZ), and along-strike strain decoupling by the CCM transfer zone. Note formation of the Agrio caldera pull-apart (AC), Copahue-Antiñir thrust front (CAF), and Lago de la Laja fault zone (LLFZ). (D) Narrowing of the volcanic arc, localization of the deformation along the axial intra-arc and Liquiñe-Ofqui fault zone. Focal mechanism is from Barrientos and Acevedo (1992). The western extent of the northern and southern Loncopué trough (NLT, SLT) (Ramos, 1977) and associated backarc volcanic centers are shown. Volcanic centers shown: AN—Antuco–Sierra Velluda, CA— Callaqui, CO—Copahue, AC—Agrio caldera, CM—Cordillera de Mandolegüe, LO—Lonquimay, LL—Llaima, SO—Sollipulli.
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in volcanic activity in the Main Cordillera and compressional deformation, resulting in shortening, uplift, and exhumation; (3) Pliocene to early Pleistocene steepening of the slab, reestablishment of the volcanic arc, extension of the orogenic structure, and transtension at the northern limit of the intra-arc Liquiñe-Ofqui fault zone; and (4) late Pleistocene to Holocene narrowing of the volcanic arc and localized extensionaltranstensional deformation in the axial intra-arc zone. ACKNOWLEDGMENTS This work was supported by the GeoForschungsZentrum Potsdam (GFZ) Southern Andes project, PIP 4162 and PICT 059 (Conicet to V.A. Ramos) and SFB (Sonderforschungbereich) 267 Collaborative Research Center “Deformation Processes in the Andes” founded by the DFG. Melnick acknowledges founding by the Deutscher Akademischer Austausch Dienst International Quality Network (IQN) at Potsdam Universität and is grateful to M. Strecker and O. Oncken for their continuous support. We would like to thank: V. Ramos, S. Kay, J. Cembrano, T. Vietor, M. Brandon, D. Sellés, P. Alvarez, and J. Clavero for fruitful discussions, and F. Charlet, M. de Batist, B. Scharf, and O. Buettner for sharing their seismic lines and bathymetry from the Lago de la Laja. Reviews by E. Godoy and A. Meigs greatly helped to improve the ideas presented in this work. REFERENCES CITED Allmendinger, R.W., Isacks, B.L., Jordan, T.E., and Kay, S.M., 1997, The evolution of the Altiplano-Puna plateau of the Central Andes: Annual Reviews of Earth Sciences, v. 25, p. 139–174, doi: 10.1146/annurev.earth.25.1.139. Angermann, D., Klotz, J., and Reiberg, C., 1999, Space-geodetic estimation of the Nazca–South American Euler vector: Earth and Planetary Sciences Letters, v. 171, p. 329–334, doi: 10.1016/S0012-821X(99)00173-9. Barrientos, S., and Acevedo, P., 1992, Seismological aspects of the 1988–1989 Lonquimay (Chile) volcanic eruption: Journal of Volcanology and Geothermal Research, v. 53, p. 73–87, doi: 10.1016/0377-0273(92)90075-O. Bird, P., 2003, An updated digital model of plate boundaries: Geochemistry, Geophysics, Geosystems, v. 4, p. 1027, doi: 10.1029/2001GC000252. Bohm, M., Lüth, S., Echtler, H., Asch, G., Bataille, K., Bruhn, C., Rietbrock, A., and Wigger, P., 2002, The Southern Andes between 36° and 40°S latitude: Seismicity and average seismic velocities: Tectonophysics, v. 356, p. 275–289, doi: 10.1016/S0040-1951(02)00399-2. Burns, W.M., Jordan, T.E., Copeland, P., and Kelley, S.A., this volume, The case for extensional tectonics in the Oligocene-Miocene Southern Andes as recorded in the 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, doi: 10.1130/2006. 2407(08). Carpinelli, A., 2000, Análisis estratigráfico, paleoambiental, estructural y modelo tectono-estratigráfico de la cuenca de Cura-Mallín, VIII y IX Región, Chile, Provincia de Neuquén, Argentina [Master’s thesis]: Concepción, Chile, Universidad de Concepción, 158 p. Cembrano, J., Schermer, E., Lavenu, A., and Sanhueza, A., 2000, Contrasting nature of deformation along an intra-arc shear zone, the Liquiñe-Ofqui fault zone, southern Chilean Andes: Tectonophysics, v. 319, p. 129–149, doi: 10.1016/S0040-1951(99)00321-2.
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Neogene tectonic evolution of the Neuquén Andes western flank Lohmar, S., 2000, Estratigrafía, petrografía y geoquímica del Volcán Antuco y sus depósitos (Andes del Sur, 37º25'S) [Master’s thesis]: Concepción, Chile, Universidad de Concepción, 118 p. López-Escobar, L., Cembrano, J., and Moreno, H., 1995, Geochemistry and tectonics of the Chilean Southern Andes basaltic Quaternary volcanism (37°–46°S): Revista Geológica de Chile, v. 22, p. 219–234. Lüth, S., and ISSA 2000 Working Group, 2003, A crustal model along 39°S from a seismic refraction profile: ISSA 2000: Revista Geológica de Chile, v. 30, p. 83–101. Melnick, D., Folguera, A., Rosenau, M., Echtler, H., and Potent, S., 2002, Tectonics from the northern segment of the Liquiñe-Ofqui fault system (37°–39°S), Patagonian Andes, in 5th International Symposium of Andean Geodynamics [extended abstracts]: Toulouse, France, IRD (Institut de Recherche pour le Développement), p. 413–417. Melnick, D., Folguera, A., Echtler, H., Charlet, F., Büttner, O., Chapron, E., Scharf, B., De Batist, M., and Vietor, T., 2003, The Lago de la Laja fault system: Active intra-arc collapse in the southern Central Andes (37°15′S): X Congreso Geológico Chileno: Concepción, Chile, Universidad de Concepción, CD-ROM. Melnick, D., Folguera, A., and Ramos, V.A., 2006, Structural control on arc volcanism: The Caviahue-Copahue complex (38°S): Journal of South American Earth Sciences (in press). Moreno, H., and Gardeweg, M., 1989, La erupción reciente del complejo volcánico Lonquimay (Diciembre de 1988), Andes del Sur: Revista Geológica de Chile, v. 16, p. 93–117. Moreno, H., and López-Escobar, L., 1994, Los centros eruptivos de Rucapillán: Actividad volcánica reciente en la Depressión Central de los Andes del Sur (39°S): VII Congreso Geológico Chileno: Concepción, Chile, Universidad de Concepción, p. 334–339. Moreno, H., Thiele, R., Lahsen, A., Varela, J., López-Escobar, L., and Vergara, M., 1985, Geocronología de rocas volcánicas Cuaternarias en los Andes del Sur entre las latitudes 37º y 38ºS, Chile: Revista Asociación Geológica Argentina, v. 40, p. 297–299. Mpodozis, C., and Ramos, V.A., 1989, The Andes of Chile and Argentina, in Ericksen, G.E., Cañas Pinochet, M.T., and Reinemund, J.A., eds., Geology of the Andes and its relation to hydrocarbon and mineral resources: Circum-Pacific Council for Energy and Mineral Resources, Houston, Texas: Earth Science Series, v. 11, p. 59–90. Muñoz, J., and Niemeyer, H., 1984, Hoja Laguna del Maule, regiones del Maule y Bío-Bío, carta geológica de Chile N°64: Santiago, Chile, Servicio Nacional de Geología y Minería, scale 1:250,000, 1 sheet, 98 p. Muñoz, J., and Stern, C., 1988, The Quaternary volcanic belt of the southern continental margin of South America: Transverse structural and petrochemical variations across the segment between 38° and 39°S: Journal of South American Earth Sciences, v. 1, p. 147–161, doi: 10.1016/08959811(88)90032-6. Muñoz, J., Troncoso, R., Duhart, P., Crignola, P., Farmer, L., and Stern, C., 2000, The relation of the mid-Tertiary coastal magmatic belt in southcentral Chile to the late Oligocene increase in plate convergence rate: Revista Geológica de Chile, v. 27, p. 177–203. Nakamura, K., 1977, Volcanoes as possible indicators of tectonic stress orientation—Principle and proposal: Journal of Volcanology and Geothermal Research, v. 2, p. 1–16, doi: 10.1016/0377-0273(77)90012-9. Naranjo, J.A., and Moreno, H., 1991, Actividad explosiva postglacial en el volcán Llaima, Andes del Sur (38°45′S): Revista Geológica de Chile, v. 18, p. 69–80. Niemeyer, H., and Muñoz, J., 1983, Hoja Laguna de la Laja, región del BíoBío, carta geológica de Chile N°57: Santiago, Chile, Servicio Nacional de Geología y Minería, scale 1:250,000, 1 sheet, 52 p. Pankhurst, R.D., Weaver, S.D., Hervé, F., and Larrondo, P., 1999, MesozoicCenozoic evolution of the northern Patagonian Batholith in Aysén, southern Chile: Journal of the Royal Society of London, v. 156, p. 673–694. Potent, S., and Reuther, C.D., 2001, Neogene deformationsprozesse im aktiven magmatischen bogen südzentral Chile zwischen 37° und 39°S: Mit-
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Geological Society of America Special Paper 407 2006
Intraplate deformation in the Neuquén Embayment Alfonso Mosquera* Tecpetrol S.A., Della Paulera 299, Piso 12, C1001ADA, Buenos Aires, Argentina Víctor A. Ramos* Laboratorio de Tectónica Andina, Universidad de Buenos Aires, Consejo Nacional de Investigaciones Científicas y Técnicas, Buenos Aires, Argentina
ABSTRACT The Neuquén Embayment, which developed along the eastern foothills of the southern Central Andes, has a complex history of intraplate deformation. The Paleozoic basement fabrics exerted a major influence in Mesozoic and Cenozoic deformation. The most important feature is an E-W–striking fault system that is related to a late Paleozoic fabric and is associated with the Huincul basement high, which truncates the basin. This fabric is interpreted as being the result of the accretion of the Patagonia terrane with Gondwana during the Early Permian. Two-dimensional (2D) and three-dimensional (3D) seismic coverage and subsurface information identify different sectors in the Neuquén Embayment that record alternating episodes of contraction and extension during the Jurassic and Cretaceous. The deformation history east of the thrust front of the Agrio fold-and-thrust belt is characterized by periods of (1) transpression and almost orthogonal contraction to the continental margin, (2) extension, and (3) relative quiescence, which alternates in different sectors. The earliest shortening occurred in the Early Jurassic when the main stress was oriented in the N-NW sector. The stress rotated to the northwest up to Valanginian times, when a more orthogonal orientation to the continental margin became dominant and prevailed after the Cenomanian. After a period of quiescence in the Neuquén Embayment associated with very oblique subduction during the Paleogene, the final contractional deformation took place in the late Miocene, with a west-east orientation of the main stress, and was followed by Pliocene extension. The changing stress patterns correlate with differences in convergence vectors between the Aluk, Farallon, and Nazca oceanic plates and the Gondwana or South American continental plates. The Aluk stage from the Jurassic to the Valanginian was characterized by tectonic inversion that is shown by shortening and right-lateral strike-slip structures that are concentrated in the Huincul system and more subtle deformation in the Chihuidos and Entre Lomas systems. The early Farallon stage was distinguished by reduced inversion and displacement in the Huincul system and a general retreat of deformation after the Valanginian. The change to late Farallon stage was characterized by a prominent tectonic inversion of the Entre Lomas system, which resulted from the inception of the formation of the Agrio fold-and-thrust *E-mails:
[email protected];
[email protected]. Mosquera, A., and Ramos, V.A., 2006, Intraplate deformation in the Neuquén Embayment, 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. 97–123, doi: 10.1130/2006.2407(05). For permission to copy, contact
[email protected]. ©2006 Geological Society of America. All rights reserved.
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A. Mosquera and V.A. Ramos belt in the retroarc area. This belt developed during most of the Late Cretaceous, when the embayment showed a general quiescence. The Nazca stage was characterized by the main episode of uplift, tectonic inversion of the older half-grabens, and important strike-slip faulting that was followed by local collapse of some structures during the Pliocene. Keywords: Neuquén Basin Embayment, Mesozoic-Cenozoic intraplate deformation, plate convergence.
INTRODUCTION The Neuquén Embayment located in the foothills of the Andes between 36°S and 41°S latitude is one of the most prolific hydrocarbon basins in southern South America (estimated ultimate recovery [EUR] of 9700 million of barrels of oil equivalent [MMBOE]). The embayment has a complex structural history, with different sectors of the Andean foreland being deformed and uplifted in an episodic way from the Early Jurassic to the late Cenozoic. The kinematics, the geometry, and the age of this intraplate deformation have been poorly known due to complex patterns of strain partitioning, different structural trends, and limited time constraints. The three-dimensional (3D) seismicreflection surveys obtained by the petroleum industry in recent years provide an excellent database to identify the orientation and ages of the different structures. The interpretation of these surveys combined with subsurface data derived from exploration and appraisal wildcat wells and extensive sequence stratigraphic studies show striking spatial and temporal patterns in this deformation. These data are evaluated herein with information derived from oceanic plate motions and kinematics, and the tectonic history of the Neuquén Andes. As a whole, the data show robust evidence for links between plate convergence directions, basement anisotropies, strain orientations in different parts of the foreland sectors, and for changes through time. Study Area The Neuquén Embayment developed east of the Principal Cordillera of Neuquén as a complex retroarc-foreland basin during Mesozoic and Cenozoic times (Figs. 1 and 2). The study area encompasses the central part of the basin east of the Miocene orogenic front and its southern and eastern boundaries. This foreland area is bounded to the northeast by the Precambrian to Paleozoic basement of the San Rafael block, which is part of the exotic Cuyania terrane, a Laurentian-derived block amalgamated to the Gondwana continent in Middle to Late Ordovician times (Fig. 3). The southeastern limit is controlled by the Patagonia terrane, a parautochthonous Gondwanan block that collided with Gondwana in Early Permian times (see Ramos, 2004a, 2004b). The foreland structural pattern of the Neuquén Embayment is strongly constrained by penetrative deformational patterns
produced in the Paleozoic (Fig. 3). This deformation controlled the orientation of the fabrics of the metamorphic basement, the faults associated with subsequent rifting during the opening of the South Atlantic Ocean (Franzese and Spalletti, 2001), and Mesozoic and Cenozoic intraplate deformation. Previous Work The present knowledge of the structure of the Neuquén Embayment is the result of several generations of geological studies. Most of them were concentrated in the southern part of the embayment along the Huincul ridge (Fig. 1). They begin with the pioneering work of Groeber (1929), Suero (1939, 1951), Herrero Ducloux (1946), and De Ferraris (1947). De Ferraris (1947) was the first to recognize the Huincul ridge or high, a prominent transverse structural feature with a W-NW to E trend that crosses the entire foreland basin. The early interpretations emphasized the extensional nature of this arch and demonstrated the role of fault control on rapid thickness changes in the different Early Jurassic units. Harding (1973, 1974) was the first to interpret the Huincul ridge as a rightlateral strike-slip system based on the en-echelon pattern of the oblique anticlines that developed along the ridge. The detailed studies by Orchuela et al. (1981), Orchuela and Ploszkiewicz (1984), and Ploszkiewicz et al. (1984) showed a combination of transpressional and transtensional structures linked with the right-lateral displacement of the Huincul ridge–bounding faults. Later, Eisner (1991) reinterpreted the Huincul ridge as a series of oblique and transverse normal faults transpressionally inverted during Andean deformation. Recent remapping and reinterpretation of outcrops by Zabala and González (2001) in the western part of the Huincul ridge, as well as 3D seismic-survey acquisition and interpretations in the central and eastern portions (Veiga et al., 2001; Mosquera, 2002; Pángaro et al., 2002; Berdini et al., 2002, Gómez Omil et al., 2002; Fernández et al., 2003) have resulted in delineation of its boundaries and the assignment of an Early Jurassic age to the beginning of strike-slip activity along the Huincul ridge. The eastern and central parts of the Neuquén Embayment have been less intensively studied due to the lack of good exposures. Most of what is known in this region is based on seismic and wildcat well data. The regional studies of Ramos (1977)
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Figure 1. Location of the study area in the Neuquén Embayment (in box) relative to the Miocene orogenic front, the Miocene intraplate deformation front, the maximum extent of the retroarc foreland basin, and the Huincul ridge.
and Digregorio and Uliana (1980) were the first to include analyses of the structure of the eastern margin of the Neuquén Embayment. These studies showed the influence of the basement fabrics in the orientation of the structures, the multiepisodic character of the deformation (Triassic to Tertiary), and a structural style characterized by low relief with gently dipping anticlines and structural noses. Orchuela and Ploszkiewicz (1984) performed a detailed study of the structure of the eastern margin of the embayment and recognized the importance of contractional Mesozoic and Cenozoic deformation in this part of the basin. They recognized the Mesozoic age of the Entre Lomas and Charco Bayo anticlines northeast of the city of Neuquén, and their reactivation during Cenozoic times. Subsequent studies have dealt with the evolution of single structures (Arregui et al., 1996). The recent acquisition of 3D seismic data, as well as the reprocessing of 2D seismic along with subsurface data, provides an excellent database to reanalyze the tectonic evolution of the northeast border of the Neuquén Embayment. The regional structure of the Neuquén Embayment shows the basement control on the main morphostructural units (Ramos,
1977). This control has been confirmed by the comprehensive tectonic evaluation of the present structures by Vergani et al. (1995). The complexity of these structures is enhanced by changes in strain orientation through time. Many other studies that have dealt with local structures are mentioned in the following. STRATIGRAPHY The Neuquén Basin sedimentary thickness totals 10,000 m and records 220 m.y. of Triassic to Tertiary basin subsidence (Fig. 2). The Mesozoic sedimentary fill is divided into seven major sequences: pre-Cuyo, Cuyo, Lotena (Jurassic), Mendoza (Jurassic-Cretaceous), Rayoso, Neuquén (Cretaceous), and Malargüe (Cretaceous-Tertiary) Groups. The oldest deposits are the pre–Cuyo Group volcanic and volcaniclastic rocks and associated clastics. They are bounded by two unconformities, the intra-Triassic at the base (215 Ma) and the intra-Liassic at the top (208 Ma). These sequences were deposited in isolated half-grabens. Their thicknesses vary from 200 m in the Medanito depocenter to over 3000 m in deep halfgrabens, like the Estancia Vieja along the Huincul ridge. These
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Figure 2. Generalized stratigraphic column of the Neuquén Basin and its relation to convergence vectors along the Pacific margin (modified from Zapata et al., 1999).
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Figure 3. Map of basement fabrics and pre-Mesozoic plate boundaries and terrane sutures in the Neuquén Basin, modified from Chernicoff and Zappetini (2003, 2004). Inset in upper right shows terrane map of Ramos (1996). The convergence vector of the oceanic Aluk plate at approximately 190 Ma and the maximum extent of foreland basin development (dashed line) near 90 Ma are indicated. The irregular dashed area shows the extent of two-dimensional (2D) and three-dimensional (3D) seismic coverage discussed in the paper. The structural fabrics indicated in the Chilenia terrane are shown in more detail in the seismic sections in Figures 5 and 6 and the time (depth) structure maps in Figure 7.
depocenters formed in a period of 7 to 10 m.y. in the Early Jurassic (Hettangian-Sinemurian). The next deposits are sedimentary sequences. The first four groups are associated with marine transgressions from the Pacific that took place before the uplift of the Andean Cordillera. The first of these is the Cuyo Group (PliensbachianBathonian), which includes the deep-marine organic-rich shales of the Los Molles Formation, the shallow-marine to deltaic deposits of the Lajas Formation, and the fluvial deposits of the Punta Rosada and Challacó Formations. The second is the Lotena Group, which consists of the Auquilco Formation evaporites, the La Manga Formation limestones, and the deltaic to shallow-marine Lotena Formation clastic deposits. The third is the Mendoza Group (late Kimmeridgian–Barremian), which is composed of the continental Tordillo Formation sandstones, the Vaca Muerta Formation organic-rich deep-marine shaly marls, the Quintuco Formation nearshore limestones, the continental to shallow-marine Mulichinco Formation and the marine clastics and limestones of the Agrio Formation. The fourth is the
Rayoso Group (Aptian-Albian), which includes the evaporitic and clastic deposits of the Huitrin and Rayoso Formations. The rest are synorogenic deposits that record the uplift of the Andes at this latitude. The Neuquén Group (CenomanianCampanian) consists of continental molasse deposits. Andean uplift started in the Late Cretaceous, as shown by fission-track dating of zircons in the lower section of the Neuquén Group (88 ± 3.9 Ma) in the Huincul Formation at Cerro Policía (see Corbella et al., 2004). The overlying Malargüe Group (MaastrichtianPaleocene) is composed of shallow-marine deposits that were deposited in association with the first Atlantic transgression. Eocene and Miocene synorogenic deposits and Pliocene volcanics complete the sedimentary record. Basement Fabrics The basement of the Neuquén Embayment is the result of the interaction of three basement terranes that were accreted to the Gondwana margin in the early to late Paleozoic (Fig. 3; see
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Ramos, 2000, 2004b). The Ordovician collision of the N-S– trending Cuyania terrane with Gondwana was followed by the Late Devonian collision of the Chilenia terrane. These two collisions produced N-NW–trending sutures. The subsequent Early Permian accretion of Patagonia produced an almost E-W– trending suture. These terrane amalgamations created complex fabrics in the basement of the Neuquén Embayment that are recognized by their magnetic patterns and residual Bouguer anomalies (Fig 4; Chernicoff and Zappetini, 2003, 2004). The Chilenia terrane, a 400-km-wide belt of continental basement (Ramos et al., 1984), underlies the largest portion of the Neuquén Basin (Figs. 3 and 4). It covers the central and western part of the Neuquén Embayment beneath the Los Chihuidos high, the Agrio fold-and-thrust belt, and the Entre Lomas system (these structural features and their location are discussed in the following). Its southern margin constitutes the central and northern part of the Huincul ridge. It has an N-NW– trending original fabric, which controls the development of the basement structures of the western and central parts of the Neuquén Embayment, the Agrio fold-and-thrust belt, and the Chihuidos anticlines. The southern margin of the terrane shows a strong E-W truncation of the fabrics that resulted from the collision of the Patagonia terrane as shown in Figure 3. The Cuyania terrane (Figs. 3 and 4) is only 100 km wide at the latitude of the Neuquén Basin in the study region. This Laurentian-derived block was amalgamated to Gondwana in the Middle to Late Ordovician (e.g., Astini et al., 1996; Ramos 2004b) and constitutes the eastern boundary of the basin. Its N-NW–trending fabrics extend 100 km west of the suture between Chilenia and Cuyania and are reflected in the Charco Bayo, Entre Lomas, and the Estancia Vieja trends (these structural features and their location are discussed in the following
section). This clearly indicates that basement strain associated with the Paleozoic suture has a broad influence. Finally, the Patagonia terrane (Fig. 3), a parautochthonous Gondwanan block that collided with Gondwana in Early Permian times (see Ramos, 2004a), defines the southern limit of the basin. The northern margin corresponds with the southern portion of the Huincul ridge and the Picún Leufú subbasin south of the ridge. The basement fabrics show mostly an E-W orientation (Figs. 3 and 4). Seismic data depict the basement fabrics in several specific places. The seismic line in Figure 5 shows an angular unconformity between the late Paleozoic metamorphic rocks of the Colohuincul Formation and the Permian to Early Triassic volcanic rocks of the Choiyoi Group in the southern edge of the basin (Fig. 3). The Colohuincul Formation generally exhibits N-dipping fold patterns with an E-W–trending fold axis. Outcrops of the Colohuincul Formation metamorphic rocks a few kilometers to the south of the seismic line show a top-to-thesouth vergence. Due to their strong ductile deformational fabrics and metamorphic grade, their age was considered to be Precambrian for many years. However, recent studies of detrital zircons in sequences of the Colohuincul Formation along the Limay River in the Bariloche region indicate a maximum age of late Carboniferous to Early Permian (Basei et al., 1999; Varela et al., 2001). The seismic lines in Figure 6 along the Huincul ridge (Fig. 3) show angular unconformities between the pre–Cuyo Group and older units that could mark the contact between the Choiyoi Group volcanics and the deformed rocks of the Colohuincul Formation. A restoration to prerift conditions shows that these basement units had a northerly dip. The Patagonia suture and its related fabrics developed in Early Permian times (von Gosen 2003). This is the most impor-
Figure 4. Residual aeromagnetic map and Bouguer residual gravimetric anomaly map with Paleozoic terranes superimposed (based on Chernicoff and Zappetini, 2003, 2004).
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Figure 5. Seismic section across the China Muerta depocenter showing the angular unconformity of the Choiyoi Group volcanic and volcaniclastic rocks with the Colohuincul Formation metamorphic rocks. Structures in the Colohuincul Formation have a dominant south vergence. Section is flattened to the bottom of the pre–Cuyo Group. Location of section is shown in Figure 3. TWT is two-way traveltime.
tant basement element that influenced the deformational history of the Neuquén Embayment. This E-W fabric is superimposed on the early Paleozoic Chilenia and Cuyania fabrics up to at least 120 km north of the Huincul ridge, as shown by the trends inferred from seismic data in Figure 3. This superposition leads to a complex pattern. The Chilenia and Cuyania terranes created fabrics and sutures with N to NW trends that were partially obliterated by late Paleozoic deformation. In some areas, pre–Choiyoi Group volcanics were affected by compressional deformation, as seen in parts of the San Rafael system (Kleiman, 2002). This deformation in the north was associated with the San Rafael orogenic episode, but further south was related to the ductile deformation related to the collision of Patagonia (Ramos, 1984, 2004a; Von Gosen, 2003). The penetrative deformation produced in the Paleozoic and the associated fabrics in the metamorphic basement played a key role during the rifting associated with the opening of the South Atlantic Ocean (Franzese and Spalletti, 2001), as well as during Mesozoic and Cenozoic backarc and foreland deformation. The Paleozoic sutures and basement fabrics were reactivated as weakness zones that controlled the extensional faults produced in association with pre–Cuyo Group synrift sedimentation (see Figs. 3, 7A, and 7B).
Rifting and Early Inversion Depocenters The initial rifting of the basin during Triassic and Early Jurassic (Sinemurian) times associated with the breakup of Pangea (Gulisano et al., 1984) formed the series of depocenters depicted in Figures 8 and 9. The pre–Cuyo Group volcanic, volcaniclastic, and clastic deposits are the synrift fill of the halfgrabens. The rift phase did not last more than 7–10 m.y. (215–208 Ma) and resulted in a dense mosaic of medium- to small-scale, closely spaced depocenters (10–70 km long × 5–10 km wide). The seismic lines in Figures 9–13 show that these depocenters are half-grabens with clearly defined wedgeshaped synrift geometries that suggest listric normal faulting. An almost-constant fault polarity creates a trend with very few conjugate faults. The pre–Cuyo Group half-grabens have three main orientations: N-S, E-W, and NE-SW (Fig. 8). The western depocenters in Figure 8, including Cordillera del Viento, Tres Chorros, Chacaico–Catán Lil, China Muerta (Fig. 12), Sañicó, Las Coloradas, and Los Chihuidos (Fig. 13), are located along the N-S trend. Their thicknesses reach up to 3000 m. They developed along the western part of the basin along the Los Chihuidos high, the Agrio fold-and-thrust belt, and the Patagonian Cordillera, and are intercepted by E-W– oriented depocenters along the Huincul ridge (Fig. 14), which
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Figure 6. Seismic section across the Las Chivas depocenter with: (A) pre–Cuyo Group half-graben flattened to the top of the fill, and (B) prerift angular unconformity flattened to the bottom of the synrift deposits. Section shows the south vergence of the basement structures in the Choiyoi Group and Colohuincul Formation(?). Location of section is shown in Figure. 3. TWT is two-way traveltime.
Figure 7. Time structure maps at the top of the pre–Cuyo Group based on three-dimensional (3D) seismic data showing fault patterns related to basement fabrics for: (A) Entre Lomas system (see Fig. 8), with W- to NW-trending patterns, and (B) Huincul ridge area, with W- to WNW-trending patterns. Location of maps is shown on Figures 3 and 4.
include Vaca Muerta–Medanito, Challacó, Bajada Colorada, Plaza Huincul, Sierra Barrosa, Río Neuquén, and El Mangrullo. These depocenters are associated with rifting along the Permian suture of the Patagonia terrane. Maximum sediment thicknesses of these depocenters reach more than 3000 m along the Huincul ridge. These depocenters form the main trend of the Neuquén Basin and have a widespread areal distribution that not only parallels the Huincul ridge, but also extends into the Añelo foredeep and the Entre Lomas broken foreland system (see Fig. 14). The third set along the NE-SW trend includes the Entre Lomas (Figs. 9 and 10), Bajada Vidal, and Estancia Vieja (Fig. 11) depocenters in Figure 8. Their northeast trend is controlled by basement fabrics that developed during the accretion
of Cuyania. Associated synrift deposits have thickness varying between 2000 m (Entre Lomas) and 3500 m (Estancia Vieja). These depocenters are located along the eastern margin of the basin where they overlap with less conspicuous E-W–trending depocenters like the 250-m-thick El Medanito depocenter. Evidence for inversion of some of these half-graben depocenters at the time of the deposition of the upper part of the pre–Cuyo Group comes from growth strata patterns and angular and erosional unconformities. Depocenters showing this pattern occur mainly along the margins of the basin, suggesting the existence of two deformation belts subparallel to the Limay and Colorado Rivers. Along the eastern margin, the seismic section in Figure 11 shows that the Estancia Vieja depocenter was
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Figure 8. Map modified from Vergani et al. (1995) showing extensional faults and major depocenters of the Neuquén Embayment. Abbreviations are: CV—Cordillera del Viento, TCh—Tres Chorros, VM—Vaca Muerta–El Medanito, CL—Catan Lil, LC—Las Coloradas–Cerro Calcatre, Sa—Sañico, CHM—China Muerta, ES—El Sauce, BC—Bajada Colorada, PH—Plaza Huincul, DCh—Dorso de los Chihuidos, SR—Sierra de Reyes, CC—Cara Cura, EL—Entre Lomas, BV—Bajada Vidal, RN—Río Neuquén, EV—Estancia Vieja, EM—El Mangrullo. Dashed lines indicate locations of seismic sections (SS1 to SS5) shown in Figures 9 to 13.
Figure 10. Seismic section SS2 across the Entre Lomas depocenter flattened to the top of the pre–Cuyo Group. Section shows the details of the intra–pre–Cuyo Group unconformity that formed at the end of the rift phase and the beginning of the inversion stage in the Entre Lomas system. Location of section is shown in Figure 8. TWT is two-way traveltime.
Figure 9. Regional seismic section SS1 from the Huincul ridge in the south (Sierra Barrosa–Los Bastos depocenters) to the Entre Lomas–Charco Bayo depocenters in the northeast. Section is flattened to the top of the pre–Cuyo Group. Location of section is shown in Figure 8. TWT is two-way traveltime.
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Figure 11. Seismic section SS3 across the Estancia Vieja depocenter flattened to the top of the pre–Cuyo Group. Section shows the details of the intra–pre–Cuyo Group angular unconformity that indicates the transition between the rift phase and the inversion stage in the eastern portion of the Huincul system. Location of section is shown in Figure 8. TWT is two-way traveltime.
Figure 12. Seismic section SS4 across the China Muerta depocenter flattened to a horizon in the intra Cuyo Group. Section shows details of the pre–Cuyo Group deformation. Location of section is shown in Figure 8.
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Figure 13. Seismic section SS4 across the Dorsal de Los Chihuidos high depocenter flattened to the top of pre–Cuyo Group. Location of section is shown in Figure 8.
mosaic of modern morphostructural units. Many of the structural features that formed in the Mesozoic experienced an important reactivation in the Miocene. In detail, the Miocene deformation front (eastern deformation limit) in the Neuquén Embayment extends up to 600 km east of the trench axis and 300 km east of the Miocene orogenic (topographic) front to the west (Fig. 1). The Miocene deformation front has a triangular shape in map view that is subparallel to the Limay and Colorado Rivers and that coincides fairly closely with the position of the Mesozoic deformation front (Fig. 14). The Miocene deformation front between 39° and 40°S latitude changes trend from N to NW to almost E-W along the Huincul ridge, where it reaches its maximum eastward extension (400 km). The bend coincides with the transition between the Patagonian Andes and the Principal Andes in the Neuquén Cordillera (Chotin and Giret, 1979). Between 39° and 37°S, the deformation front changes orientation from E-W to NW following the trend of the Entre Lomas fault system (Figs. 1 and 8). These bends and the eastward shift in the deformation front are unique along the Principal Andes, which everywhere else is characterized by a N-W orientation. MAIN MORPHOSTRUCTURAL UNITS
strongly inverted during the deposition of the upper pre–Cuyo Group. South of the Huincul ridge, the seismic section in Figure 12 across the China Muerta depocenter shows evidence for a strong compressional deformation during the deposition of the upper pre–Cuyo Group, resulting in partial structural inversion and segmentation of the depocenter. Similarly, the seismic line in Figure 10 provides evidence for partial inversion of the Entre Lomas system depocenters during the sedimentation of the pre–Cuyo Group. The inversion and segmentation of these depocenters seem to be the result of transpressive deformation. In contrast, no significant evidence for structural inversion is recorded during the deposition of the pre–Cuyo Group in the depocenters along the Huincul system. Inversion of these depocenters started after the deposition of the pre–Cuyo Group and before the marine transgression that produced the Cuyo Group Los Molles Formation (Pliensbachian). Top lap patterns of the upper beds of pre–Cuyo Group against the Los Molles Formation of the Cuyo Group provide clear evidence for the beginning of this inversion (see Fig. 12). These inversions mark the beginning of a larger multiepisodic intraplate deformation process that continued through the rest of the Mesozoic and into part of the Cenozoic. This deformation process was strongly influenced by the Late Triassic–Early Jurassic half-grabens. INTRAPLATE DEFORMATION Intraplate deformation in the Neuquén Embayment has been active essentially without interruption since the Early Jurassic (Hettangian-Sinemurian). This deformation, which has varied in intensity and location, has resulted in a complex
The present morphostructural configuration of the Neuquén Embayment is mainly the result of the Miocene reactivation. Seven morphostructural units have been identified and classified into three groups: the fold-and-thrust belt, the foreland basin, and the Huincul system (Fig. 14). The fold-andthrust belt group includes the Loncopué trough, the reactivated Agrio fold-and-thrust belt, the Patagonian Cordillera, and the Collón Cura basin. The foreland basin units include both the undeformed foreland, such as the stable Patagonian platform, and regions of broken foreland basin (Los Chihuidos, Entre Lomas, Picún Leufú systems, and Añelo foredeep). The Huincul intraplate system intersects and runs through both the fold-andthrust belt and the foreland basin groups. The emphasis herein is on the tectonic evolution of the Huincul system and the foreland morphostructural unit, and is based primarily on 2D and 3D seismic data, wildcat well information, and surficial geological maps. The seismic database consists of 100,000 km of 2D and 8000 km2 of 3D data and integrates interpreted data from the Tecpetrol company with published data. A regional depth structure map on the base of the Vaca Muerta Formation in Figure 15 illustrates the complexity and present-day configuration of the Huincul system and the foreland basin morphostructural units. Each morphostructural unit is analyzed herein in terms of its present configuration, tectonic evolution during the Mesozoic and Cenozoic, and the role of pre-existing structural fabrics. The morphostructural units are discussed relative to six extensional and compressional stages: (1) Early Jurassic (Hettangian-Sinemurian?) extensional stage, (2) Jurassic–Early Cretaceous (SinemurianValanginian) compressional stage, (3) Early Cretaceous
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Figure 14. Map of major morphostructural units of the Principal Cordillera and the Neuquén Embayment. The main structures in the Neuquén Embayment are based on seismic data and field evidence. Those in the fold-and-thrust belt are based on Herrero Ducloux (1946) and subsequent papers. Dashed lines labeled SS1 and SS2 are locations of seismic lines in Figures 16 and 17. The heavy dashed line shows the border of the Neuquén Embayment.
(Valanginian-Albian) compressional stage, (4) Late Cretaceous (Cenomanian-Maastrichtian) compressional stage, (5) Miocene compressional stage, and (6) Pliocene extensional stage. Fold-and-Thrust Belt Group The morphostructural units in the fold-and-thrust belt group have a simpler deformational history when compared to the multi-episodic deformation recorded in the other groups. Their tectonic evolution extends from the Late Cretaceous to the Cenozoic with the earliest contractional deformation in the Agrio fold-and-thrust belt occurring in the Late Cretaceous. In contrast, pre–100 Ma tectonic activity is recorded in the morphostructural units of the foreland. Early Jurassic half-grabens controlled the vergence of the thrusting in both regions.
Loncopué Trough. This late Cenozoic extensional trough runs parallel to the Principal Cordillera through the foothills of the Andes and the foreland. The trough is formed by a complex half-graben system that was produced during the Oligocene and extensionally reactivated in the Pliocene and Pleistocene. As a result, the Neuquén Andes between 36°30'S and 39°00'S show well-developed extensional features (Ramos, 1977; Jordan et al., 2001; Ramos and Folguera, 2005; Burns et al., this volume, chapter 8) when compared with other parts of the Central Andes where contractional stresses have been dominant. Agrio Fold-and-Thrust Belt. The Agrio fold-and-thrust belt, defined by Bracaccini (1970) and Ramos (1977), is a fossil antithetic belt with a long and complex history of deformation. The contraction and uplift of this belt is known to have begun by ca. 100 Ma and to have continued episodically through most
Figure 15. Depth structure map to the base of the Vaca Muerta Formation based on seismic data and wildcat well information. Areas with available in-house and published 2D and 3D seismic data are shown in light areas. Dashed lines labeled SS1, SS2, and SS3 are locations of seismic lines in Figures 16, 17, and 11. Contours in m.
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Intraplate deformation in the Neuquén Embayment of the Late Cretaceous (see Zamora Valcarce et al., this volume, chapter 6). Deformation terminated in the late Miocene, and there is no neotectonic activity. The deformation style is one of thin-skinned detachment folding with basement inversion in the last stages (Zapata and Folguera, 2005; Zapata et al., 1999). Pre–100 Ma contraction (Early Jurassic–Early Cretaceous) could have occurred in the Agrio fold-and-thrust belt area in association with localized transpression on some pre–Cuyo Group N-trending normal faults. Possible evidence comes from seismic data along the adjacent Los Chihuidos fault. Patagonian Cordillera. South of 38°S, the main axis of the Andes is characterized by a series of exhumed longitudinal batholiths in the Patagonian Cordillera. This region has a relatively simple deformational history. N-trending structures dominate the region. The external (western) part of the Patagonian Cordillera has a series of N-trending half-grabens, such as the Las Coloradas and Cerro Calcatre (see Fig. 8), that were inverted with a west vergence during the Miocene. No significant Mesozoic inversion is recorded in the area. Collón Cura Basin. The Collón Cura basin as defined by González Díaz and Nullo (1980) is characterized by Oligocene– early Miocene continental sedimentary rocks deposited along the foothills of the Andes during an important period of extension. This basin shares some common characteristics with the Loncopué trough, but without the evidence of recent extensional tectonics.
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Foreland Basin Group Foreland and Broken Foreland Basin. The foreland basin extends up to 400 km east of the Agrio fold-and-thrust belt and the Patagonian Cordillera (Figs. 1 and 14). North of the Huincul system, a broken foreland basin has developed, whereas to the south both a broken and an undeformed foreland coexist. The northern broken foreland includes the Chihuidos and Entre Lomas systems, which are separated by the Añelo Andean foredeep. This subdivision is based on differences in structural geometries and orientations that reflect distinct tectonic histories. This region is characterized by a series of basementinvolved thrusts (Fig. 16) that were active during the Mesozoic and then strongly reactivated during the Miocene, when the Late Cretaceous Neuquén Group was folded. The southern broken foreland south of the Huincal ridge formed in response to a less intense Miocene inversion across the Picún Leufú system. An undeformed foreland basin developed along the Patagonian Platform. Los Chihuidos–Entre Lomas Broken Foreland Basin. The Los Chihuidos and Entre Lomas fault systems played a key role in the tectonic history of the Neuquén Embayment. Both systems were characterized by continuous tectonic activity during the Jurassic and Cretaceous, but with different local stress regimes that resulted in distinct uplift histories. These fault systems are separated by the Añelo Andean foredeep, which was
Figure 16. Regional seismic section SS1 across the easternmost portion of the Agrio fold-and-thrust belt (FTB) and the Chihuidos–Entre Lomas broken foreland basin. Location of section is shown in Figures 14 and 15. TWT is two-way traveltime.
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subjected to alternating periods of active deformation and quiescence during the Mesozoic. Los Chihuidos System. The 200-km-long and 150-kmwide Los Chihuidos system is one of the Neuquén Basin’s most prominent features (Fig. 14). The main structure is a basement cored anticline with a steep western forelimb and a gently dipping eastern back limb (Fig. 16). The anticline has a gentle southern plunge toward the Huincul ridge and a steep northern plunge. The northern edge is truncated by the Cortaderas strikeslip fault (Ramos, 1981), whereas the southern edge is bounded by the E-W–trending El Mangrullo–Sauzal Bonito structure, which parallels the northern side of the Huincul ridge (Fig. 14). Seismic data in Figures 13 and 16 reveal a N-S–striking pre–Cuyo Group half-graben configuration in the roots of the Los Chihuidos anticline. Surficial geologic maps show a series of secondary folds along the axis of the main anticline that can be seen in the seismic line in Figure 16. Gravimetric data show strong positive Bouguer anomalies beneath the anticline that indicate either substantial crustal thinning or injection of mafic material into the crust (Ramos et al., 2002). Both alternatives are consistent with crustal attenuation during Late Triassic– Jurassic rifting (Vergani et al., 1995). The contractional history of the Los Chihuidos region began in the Jurassic to Early Cretaceous (SinemurianValanginian) when the Chihuidos half-grabens were inverted by transpression. A left-lateral movement along the Chihuidos fault system is inferred from the regional N-S orientation of the faults. This stage was dominated by the localized inversion of halfgrabens along the Chihuidos fault. Strong bed thinning, erosion, and local wedge geometries are clearly seen on seismic lines not presented in this paper, especially in the PliensbachianToarcian Cuyo Group Los Molles Formation. In contrast, the Early Cretaceous sequences have more tabular bed geometries and show local saddle bed thinning toward the Chihuidos fault. The Los Chihuidos proto-anticline, which formed at this time, has a symmetric shape (30 km) and extends for 80 km in an almost N to N25°W orientation. Inversion along the Chihuidos fault was severe enough to control the subsidence rate and bed thickness in the western part of the embayment. The Late Cretaceous stage cannot be analyzed on available seismic lines. However, outcrop data show a low-angle unconformity (interCenomanian) between the upper beds of the Rayoso Group and the Neuquén Group on the forelimb of the Los Chihuidos anticline, indicating a reactivation pulse (Herrero Ducloux, 1946; Ramos, 1981). The present configuration of the Los Chihuidos anticline is the result of Miocene inversion, which produced a major displacement along the main western fault. Other faults were only slightly inverted and produced minor structures. A small amount of Pliocene extensional reactivation of faults inverted in the Miocene is consistent with the eruption of the small intraplate alkali cones like the Parva Negra volcano (Ramos and Barbieri, 1989). Entre Lomas System. The Entre Lomas system (Figs. 8 and 14) consists of a series of NW-trending symmetric anti-
clines (Entre Lomas, Charco Bayo, Loma Montosa, and La Jarilla) and associated E-W–trending structures (Medanito, Aguada de los Indios, Puesto Morales Norte). The system extends 200 km from north to south and is 100 km across. The Estancia Vieja trend associated with the Huincul system is located at the southern termination. The eastern limit of the Entre Lomas system extends east of the Colorado River beyond the seismic coverage and the basin border. There, the Entre Lomas system extends into the Lihuel Calel hills. The northern end of the Entre Lomas system is covered by the Auca Mahuida volcanic flows, and the western limit coincides with the western flank of the Entre Lomas anticline (Fig. 14). The anticlines of the Entre Lomas system are narrow, elongated, and symmetric (5–15 km wide and 30–70 km long), have low relief (100 m), and are related to basement-involved faults. Their present geometries reflect late Miocene deformation. The eastern limit of the region affected by Mesozoic and Miocene deformation is considered to be east of the Neuquén River in the Lihuel Calel hills where basin basement outcrops. The tectonic history of the area is strongly controlled by two sets of basement anisotropies: a main one with a NW trend, and a second one that is oriented almost E-W. As in the Los Chihuidos system, contractional deformation on the Entre Lomas system began with transpressional inversion and segmentation of pre–Cuyo Group depocenters during the deposition of the upper pre–Cuyo Group (Sinemurian?). Overall, the Jurassic stage was characterized by uplift of the western flank of the Entre Lomas system and minor tectonic inversion in the central and eastern portions. The deformed region, which encompasses an area about 150 km wide, remained the same in the Early Cretaceous, but the intensity of deformation increased. The uplift of the western flank of the Entre Lomas system is reflected by the wedge geometry of the Cuyo, Lotena, and Mendoza Groups. A 2000 m reduction in sediment thickness in a distance of 20 km occurred due to onlap and truncation. The central and eastern portions of the area have a more condensed and partial sedimentary record and show minor inversions of half-grabens. E-W–oriented depocenters, like the Medanito half-graben, appear to be the most inverted during this stage. Based on crosscutting relations and deformational geometries, the Entre Lomas system is interpreted as a left-lateral strike-slip system in which half-grabens were inverted by transpression. The late Early Cretaceous stage, especially the AlbianCenomanian, was characterized by a more pronounced and generalized inversion of half-graben systems than the Jurassic. Narrow and symmetric anticlines as well as structural noses formed, and NW-oriented half-grabens were strongly inverted, creating local unconformities and growth strata patterns. This topography was peneplained by the inter-Cenomanian unconformity. The Late Cretaceous stage was characterized by a near tectonic quiescence, with deformation limited to a few gentle structures. Miocene deformation then strongly reactivated the inversion process and resulted in the folding of the inter-
Intraplate deformation in the Neuquén Embayment Cenomanian unconformity and the Neuquén Group. During the Pliocene, many structures like Entre Lomas and Loma Montosa systems experienced several hundred meters of extensional collapse (Arregui et al., 1996). Añelo Foredeep. The Añelo depocenter corresponds to the syncline between the main Los Chihuidos and Entre Lomas anticlines (Figs. 14–17). The syncline is a Miocene feature that has a NW trend that coincides with the western flank of the Entre Lomas anticline (Fig. 16). The syncline is truncated by the Cortaderas strike-slip fault (see Ramos et al, this volume, chapter 1) to the north and separates the southern part of the Entre Lomas system from the Huincul ridge to the south. Total sediment thicknesses along the Añelo foredeep reach more than 5000 m. The inversion history of the Añelo foredeep began as Early Jurassic rifting ceased during the deposition of the upper part of the pre–Cuyo Group and the rifts were either inverted or segmented by transpression. During the Jurassic to Early Cretaceous (Sinemurian-Valanginian), the inversion process continued along narrow right-lateral WNW transpressive zones and then ceased in the Valanginian. Evidence for intermittent tectonic activity comes from the thicknesses of the Cuyo, Lotena, and Mendoza Groups. These units show a marked thinning to the east due to activity on the Entre Lomas system and a gentle pinch-out to the west due to activity on the Los Chihuidos high. Growth strata patterns with wedges and truncations are clearly seen on seismic lines. The Early Cretaceous stage was characterized by a tectonic quiescence that is reflected by tabular geometries in the upper Mendoza and Rayoso Groups. The base of the Late Cretaceous Neuquén Group is marked by the inter-Cenomanian unconformity produced by the erosion of the upper Rayoso Group. A low-relief angular unconformity formed due to the tilt of the eastern part of the Añelo Andean foredeep extends eastward into the Entre Lomas system, where most of the Early Cretaceous deposits have been removed. The Late Cretaceous was characterized by a tectonic quiescence that is reflected in the tabular geometries of the Neuquén Group. Miocene deformation produced a very subtle reactivation of the rift faults and caused bending of the interCenomanian unconformity. Picún Leufú Broken Foreland Basin. This broken foreland basin lies south of the Huincul system and east of the Patagonian Cordillera along the Limay River (Fig. 14). The general shape is that of a broad, gentle syncline. In the western portion, the trend of the syncline is generally northward, correlative to the axis of the China Muerta depocenter (Fig. 8). The western flank is at the eastern foothills of the Patagonian Cordillera, whereas the eastern flank coincides with reactivated pre–Cuyo Group normal faults. In the eastern portion, the syncline strikes almost W-E. Its northern flank is formed by the southern edge of the Huincul system and its southern flank by the gentle dip of the North Patagonian massif. The Picún Leufú region is an area of low strain compared to the Los Chihuidos– Entre Lomas broken foreland system. The Picún Leufú region was characterized by a general tectonic quiescence in the Mesozoic and Tertiary, with the excep-
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tion of an intense Early Jurassic rifting and early inversion phase that particularly affected the China Muerta depocenter. Younger deformation is largely restricted to very subtle contraction deformation concentrated along a zone of weakness called the Limay lineament. The deformation front follows the Limay River. Miocene deformation caused a minor inversion in some half-grabens and again seems to be concentrated along the Limay River, particularly in the China Muerta depocenter. The deformation front is outside of the seismic coverage, but is again considered to follow the Limay River. Surficial geological data do not indicate any Mesozoic or Cenozoic formation farther south. Patagonian Platform. The Patagonian Platform includes the region of the Neuquén Basin south of the Limay River. The platform extends 300 km to the east and has a north to south width of up to 100 km. Surface geologic data show no evidence for Mesozoic or Miocene deformation, as the region is characterized by horizontal to subhorizontal Mesozoic to Tertiary beds that onlap the igneous and metamorphic rocks of the Somún Cura Massif to the south. Huincul System The Huincul ridge is the most outstanding feature of the Neuquén Embayment with regard to intensity, surficial extent, and duration of tectonic activity, which has lasted virtually without interruption from the Jurassic to the Tertiary. As shown in Figure 14, the Huincul system runs from the city of Zapala in the west to the city of Neuquén in the east. The system is actually a segment of a longer transverse feature that extends to the west into the Coastal and Principal Cordilleras of Chile (Chotin, 1976; Chotin and Giret, 1979) and to the east along the Colorado and Negro Rivers out of the seismic coverage (Fig. 8). The north-south width of the Huincul system is 120 km in the west and narrows to 60 km in the east. The system marks the boundary between the broken foreland basins and the Añelo foredeep to the north and the Picún Leufú region to the south (Fig. 14). The central part of the Huincul region is characterized by a series of reactivated normal faults that coincide with the Huincul ridge (Fig. 17). This region is characterized by half-graben faults that are orthogonal to the Andean trend and that formed in a rightlateral transpressive system. These normal faults were reactivated by right-lateral strike-slip movement and reverse faults associated with narrow and locally strike-slip–controlled oblique basins (Ploszkiewicz et al., 1984). This intraplate deformational belt is considered to mark the northern boundary of the Paleozoic Patagonia terrane. Tectonic activity in the Huincul system reached a maximum in intensity and extent during the Jurassic stage, when it affected an area of 30,000 km2 (300 km E-W × 100 km N-S). The intensity of deformation diminished in the Late Cretaceous, when the affected region was reduced to 10,000 km2 (200 km length × 65 km width). This trend of decreasing deformation intensity continued into the Late Cretaceous and Tertiary.
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Figure 17. Regional seismic SS2 across (A) the Huincul ridge (south to north) and (B) the Entre Lomas broken foreland basin (southwest to northeast). Location of section is shown in Figures 14 and 15. TWT is two-way traveltime.
Intraplate deformation in the Neuquén Embayment The tectonic evolution of the Huincul ridge began with widespread Early Jurassic extension (Hettangian-Sinemurian?), which produced a series of W-WNW–oriented half-grabens with a persistent north vergence. The deepest depocenters of the Neuquén Embayment, such as the Estancia Vieja and Los Bastos depocenters, formed along the Huincul ridge. The subsequent stage was characterized by the inversion of these half-grabens at the time of deposition of the upper pre–Cuyo Group. This inversion was more intense in depocenters like the Estancia Vieja along the eastern end of the Huincul ridge than in depocenters like los Bastos, Pampa Bandera, and Sierra Barrosa in the central and western portions. Jurassic deformation coincided with the maximum expansion of the Huincul system, as half-graben faults were reactivated by reverse and right-lateral strike-slip faulting. Narrow, local strike-slip basins also formed. The associated anticlines are characterized by steep S-dipping forelimbs and gentle N-dipping back-limb geometries, suggesting that the main displacement component was dip-slip in half-grabens like Sierra Barrosa and Los Bastos. Compression on right-lateral strike-slip displacement appears to have been a secondary process. Growth strata geome-
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tries, like wedges, onlaps, and truncations, are common in the pre–Cuyo Group and are especially common in the Cuyo, Lotena, and Lower Mendoza Groups along the Huincul system. Extensional to transtensional small faults developed in most of the anticlines, especially in the crestal positions of the lower part of the Mendoza Group (Vaca Muerta and Quintuco Formations). This extensional episode had a widespread distribution along the Huincul system, and its duration is very well constrained, coinciding with the end of the Jurassic-Valanginian stage. The Early Cretaceous stage was characterized by a reduction of about 30% in the areal extent of the Huincul system (Figs. 18 and 19). This reduction was the result of a southward and westward retreat of the Jurassic compressional deformation front (compare Figs. 18 and 20). The retreat took place along the eastern end of the Huincul system where Jurassic-stage anticlines, like those in the Estancia Vieja trend, collapsed by extensional-transtensional faulting. Tectonic activity continued in the eastern end of the Huincul system, but was restrained to a 10-km-wide zone around the main Huincul fault shown in the easternmost part of Figure 20.
Figure 18. Structures developed during the Aluk stage between the Sinemurian and Valanginian. Shaded area indicates region of maximum strain during this period. Dashed faults correspond to incipient displacements.
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Figure 19. Regional seismic section SS1 across the Chihuidos–Añelo foredeep–Entre Lomas systems of the broken foreland basin in Figure 16 flattened to the Valanginian. Location of section is shown in Figures 8, 14, 15, and 18. TWT is two-way traveltime.
Large-scale collapse of anticlines also took place in the western portion of the Huincul ridge in the Ramon Lista extensional system (see Pángaro and Bruveris, 1990). These authors described active extensional faulting during the Early Cretaceous. Other extensional faults are present, but at local scale and related to the lateral displacement of specific faults. A 40 km southward retreat of the compressional deformational front took place along the central part of the Huincul system. In that region, the northernmost half-grabens ceased inverting, and tectonic activity became concentrated along a 40-km-wide zone. The Early Cretaceous beds of the Agrio Formation and Rayoso Group show an almost tabular geometry in the external portions of the Huincul system and clear growth strata patterns along the central part. Deformation during the Late Cretaceous was constrained to an even narrower zone along the main Huincul fault system. Miocene deformation resulted in reactivation of inversion that led to folding of the Neuquén Group beds and the interCenomanian unconformity. The Pliocene stage resulted in collapse of the system along the main Huincul faults.
et al. (1990), and Scheuber et al. (1994). These authors recognized important changes in the orientation of the convergence vectors in the subduction zone between the adjacent oceanic plates and the South American plate (or western Gondwana plate). As early as the beginning of the separation between the North and South American plates and the formation of the Caribbean plate at about 160 Ma, the relative convergence vector between the Aluk and South America plates had a strong N-NW oblique direction relative to the continental margin. The vector between the Farallon and South American plates then rotated from a NW orientation to a more orthogonal position in the Late Cretaceous (Zonenshayn et al., 1984, 1987). A period of more oblique convergence followed in the Paleogene, when the convergence vector between the Farallon and South America plates was oriented to the SW. The convergence direction became more orthogonal at the latitude of Neuquén with the breakup of the Farallon plate and the beginning of the subduction of the Nazca plate at about 26–27 Ma. The present convergence direction is slightly oblique to the margin with an orientation of N105°W.
TECTONIC INTERPRETATION
Main Oceanic and Foreland Stages of Intraplate Deformation
In order to interpret the intraplate deformation in sectors of the Neuquén Embayment, we analyzed the oceanic plate kinematics as described by Zonenshayn et al. (1984, 1987), Jaillard
The Mesozoic and Tertiary tectonic evolution of the Neuquén Embayment can be subdivided into a series of deformational stages related to strain orientations in the different sectors.
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Figure 20. Structures developed during the early Farallon stage between the Valanginian and Cenomanian. Shaded area indicates region of maximum strain during this stage. Dashed faults correspond to inactive or incipient displacements.
Based on this analysis, the Aluk, early Farallon, late Farallon, and Nazca stages of deformation are recognized and discussed in the following sections. Aluk Stage. In the Early Jurassic, deformation was largely concentrated in the Huincul ridge, as reflected by the inversion and right-lateral slip displacement of W- and W-NW–trending half-grabens (Fig. 18). Contemporaneous deformation in the Los Chihuidos and Entre Lomas systems was more subtle (see Fig. 19), as reflected by the formation of narrow and lowamplitude anticlines related to left-lateral displacement zones. Minor growth strata developed. As seen in Figure 18, the structures along the Huincul system show a general orientation of N42–45°W for the main strain across the region at this time. This orientation fits well with the Jurassic oceanic plate kinematic picture of Zonenshayn et al. (1984, 1987) and Jaillard et al. (1990), since the oceanic convergence vector had a N to N-NW orientation at this time. This vector originated at about 160 Ma during the early separation of the North and South American plates by spreading related to the formation of the Caribbean plate. The very strong oblique convergence of the Aluk plate with the continental margin of Neuquén explains a near lack of strain partitioning in the foreland. This configura-
tion can explain why the Huincul system, which was relatively orthogonal to the main stress, concentrated most of the deformation at this time. Early Farallon Stage. As shown by the oceanic plate reconstruction of Zonenshayn et al. (1984, 1987), the convergence vector between the Pacific plates and South America was different in the early Farallon stage (Fig. 20). Following Scheuber et al. (1994) and Jaillard et al. (1990), the convergence vector rotated from N40°W to almost orthogonal to the margin during the mid-Cretaceous. The time of change corresponds to an important reduction in deformation on the Huincul ridge and a north-to-south retreat of the deformation front in the central part of the region and an east-to-west retreat in the eastern part of the region (compare Figs. 18 and 20). The retreat of the deformation front was accompanied by extensional collapses that started on both extremes of the Huincul system in the Valanginian. There is also evidence for extension on the Estancia Vieja and associated structures in the east (Fig. 21) and on the Ramon Lista area to the west, as shown by Pángaro et al. (2002). After the Valanginian, deformation continued, but lessened on the Huincul ridge and became more concentrated on the Entre Lomas system (Fig. 22). This change in deformation
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Figure 21. Seismic section SS3 across the Estancia Vieja depocenter. Location of section is shown in Figures 8, 15, 18, and 20. TWT is two-way traveltime.
Figure 22. Regional seismic section SS1 across the Chihuidos–Añelo foredeep–Entre Lomas systems of the broken foreland basin flattened to the base of the Cenomanian. Location of section is shown in Figures 8, 14, 15, 18, and 20. TWT is two-way traveltime.
Intraplate deformation in the Neuquén Embayment trend becomes even more conspicuous in the rest of the Early Cretaceous, with uplift climaxing in inter-Cenomanian times. The end of the early Farallon stage marks the beginning of shortening in the Agrio fold-and-thrust belt. Late Farallon Stage. This stage coincides with the development of the Agrio fold-and-thrust belt during the Late Cretaceous (Zamora Valcarce et al., this volume, chapter 6). Uplift of the entire Entre Lomas system is recorded by the truncation of the Rayoso Group deposits (Fig. 22). This uplift is interpreted as the potential onset of the peripheral bulge of the Agrio foldand-thrust belt along the Entre Lomas sector. To the east, most of the inner retroarc area of the Agrio fold-and-thrust belt was uplifted, and the thrust front coincided with the trace of the present Neuquén River. Outside of the Agrio belt, most of the sediments in the Neuquén Embayment region have a tabular geometry, and there is only very minor evidence for deformation. Within the basin, there is very minor evidence for Paleogene deformation of the Huantraico area north of the Los Chihuidos ridge (Cobbold and Rossello, 2003) and in the vicinity
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of the Sierra Negra, east of the Los Chihuidos area (Zapata et al., 2003). Most of the embayment records neither uplift nor sedimentation during this time. Nazca Stage. The breakup of the Farallon plate into the Cocos and Nazca plates at about 26–27 Ma coincided with a new subduction regime as convergence became more orthogonal along the Andean continental margin at these latitudes (Pardo Casas and Molnar, 1987; Somoza, 1998). The resulting contractional Miocene deformation caused the uplift of the Los Chihuidos high (Figs. 14 and 15), the final inversion of the central part of the Huincul system (Figs. 22, 23, and 24), and the final inversion that led to folding of the entire Entre Lomas area (Fig. 16). The change in convergence could also have been responsible for important right-lateral strike-slip faulting. Miocene deformation propagated into the easternmost foreland, uplifting the Lihuel Calel hills east of the Colorado River by basement back thrusting. Shortly after late Miocene deformation, widespread extension occurred across all of the Neuquén Embayment, as seen in the Huincul, Entre Lomas, and Los Chihuidos areas.
Figure 23. Seismic section across the Huincul ridge showing the final deformation of the Neuquén Group produced after Miocene compression. Location of section is shown in gray rectangle at the southern end of line labeled SS2 in Figure 14.
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Figure 24. Regional structural interpretation of the Huincul system showing the end stage of deformation after the Miocene contractional stage. Figure is based on Fernández et al. (2003).
CONCLUSIONS Our analysis of the deformation through time of the Neuquén Embayment shows that there is striking evidence for intraplate strain east of the thrust front of the Agrio fold-andthrust belt. The orientation of the strain and the timing of deformation lead to four main observations (see summary in Fig. 25): 1. The Neuquén Embayment is truncated by a major fault system known as the Huincul ridge. The basement anisotropies; the angular unconformities between the Mesozoic sequences and the Choiyoi Group volcanics and the metamorphic rocks of Colohuincul Formation; the vergence of the Paleozoic deformation; and the E-W–trending structural pattern correlate with geologic evidence in the North Patagonian massif that indicates that these anisotropies are a consequence of the collision of the Patagonia terrane with the Gondwana margin in Permian times. This major structure controlled intraplate deformation during most of the Mesozoic and Cenozoic. 2. The sequence and location of uplifts, tectonic inversion of the half-graben systems, sense and location of strike-slip displacements, magnitudes of the structures, and areas of quiescence in the Neuquén Embayment require a changing pattern of the main stress. Based on this pattern, the main stress vector is proposed to have been located in the N-NW to NW quadrant in the Early Jurassic and then to have rotated to a more orthogonal orientation in the Early to Late Cretaceous. The final uplift of the region in the Miocene can be related to an almost orthogonal compression followed by Pliocene extension in the foreland.
3. There is a striking match between the main stress orientation deduced from the structural pattern of the different sectors of the Neuquén Embayment and the convergence vectors inferred from oceanic plate kinematics. This implies that strain partitioning was low or nonexistent, and that intraplate deformation is mainly controlled by the convergence vector. The crustal erosion by subduction was minimal at these latitudes, which is consistent with little change in the orientation of the continental margin since the Paleozoic. 4. The structural patterns and the timing of deformation can be explained by four distinct convergence stages: the Aluk, from Early Jurassic to Valanginian times (Fig. 25A); the early Farallon, from Valanginian to Cenomanian time; the late Farallon, from the Cenomanian to the Paleogene (Fig. 25B); and the Nazca stage, since the Miocene. The interaction among the different oceanic plates and their relation to distinctive convergence vectors provide a comprehensive illustration of the intraplate deformation of the Neuquén Embayment.
Figure 25. Maps showing the schematic evolution of the Neuquén Embayment. (A) Active structures and convergence vector when the Aluk plate was subducting in the Sinemurian to Valanginian. (B) Active structures and convergence vectors when the Farallon plate was subducting. The early Farallon stage occurred between the Valanginian and Cenomanian, and the late Farallon stage, during the Late Cretaceous and Paleogene. Lighter gray faults show incipient displacements. Fault systems are: (1) Agrio fold-and-thrust belt, (2) Chihuidos system, (3) Entre Lomas system, and (4) Huincul system.
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ACKNOWLEDGMENTS The authors want to acknowledge many colleagues for cooperation in supplying a complete record of seismic lines and subsurface information. We especially want to thank Andres Boll from Tecpetrol, who encouraged the present study. A grant from Agencia Nacional de Promoción Científica y Tecnológica (PICT 06729/99) provided partial support for this study. Reviews by Gloria Eisenstadt and Richard Allmendinger contributed to a noteworthy improvement to an early version of the manuscript. REFERENCES CITED Arregui, C., Benotti, S., and Carbone, O., 1996, Sistemas petroleros asociados en el yacimiento Entre Lomas, Provincia del Neuquén, in 13 Congreso Geológico Argentino y 2 Congreso de Exploración y Desarrollo de Hidrocarburos: Actas, v. 1, p. 287–306. Astini, R., Ramos, V.A., Benedetto, J.L., and Vaccari, N.E., 1996, La Precordillera: Un terreno exótico a Gondwana, in XIII Congreso Geológico Argentino and III Congreso Exploración de Hidrocarburos, (Buenos Aires): Actas, v. 5, p. 293–324. Basei, M.A.S., Brito Neves, B.B., Varela, R., Teixeira, W., Siga, O., Jr., Sato, A.M., and Cingolani, C. 1999, Isotopic dating on the crystalline basement rocks of the Bariloche region, Río Negro, Argentina, in II South American Symposium on Isotope Geology: Anales Segemar v. 34, p. 15–18. Berdini, O., Arregui, C., and Pimentel Mendes, M., 2002, Evolución tectosedimantaria de la estructura Río Neuquén, Cuenca Neuquina, República Argentina, in 15 Congreso Geológico Argentino (Calafate): Actas, v. 3, p. 187–192. Bracaccini, O., 1970, Rasgos tectónicos de las acumulaciones Mesozoicas en las provincias de Mendoza y Neuquén, República Argentina: Asociación Geológica Argentina, Revista, v. 35, no. 2, p. 275–284. Burns, W.M., Jordan, T.E., Copeland, P., and Kelley, S.A., this volume, The case for extensional tectonics in the Oligocene-Miocene Southern Andes as recorded in the 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, doi: 10.1130/2006.2407(08). Chernicoff, C.J., and Zappetini, E.O., 2003, Delimitación de los terrenos tectono-estratigráficos de la región centro-austral Argentina: Evidencias aeromagnéticas: Revista Geológica de Chile, v. 30, no. 2, p. 299–316. Chernicoff, C.J., and Zapettini, E.O., 2004, Geophysical evidence for terrane boundaries in south-central Argentina: Gondwana Research, v. 7, no. 4, p. 1105–1117, doi: 10.1016/S1342-937X(05)71087-X. Chotin, P., 1976, Etude d'un secteur des Andes Meridionales (LonquimayNeuquén) a l'aide de documents satéllites ERTS-1, in 1 Congreso Geológico Chileno: Santiago, Chile, Actas, v. 1, p. B29–38. Chotin, P., and Giret, A., 1979, Analysis of northern Patagonian transverse structure (Chile, Argentina: 38° to 42° S.L.) from Landsat documents, in 7 Congreso Geológico Argentino (Neuquén): Actas, v. 2, p. 197–202. Cobbold, P.R., and Rossello, E.A., 2003, Aptian to Recent compressional deformation of the Neuquén Basin, Argentina: Marine and Petroleum Geology, v. 20, no. 5, p. 429–443, doi: 10.1016/S0264-8172(03)00077-1. Corbella, H., Novas, F.E., Apesteguía, S., and Leanza, H.A., 2004, First fission track-age for the dinosaur-bearing Neuquén Group (Upper Cretaceous) Neuquén Basin, Argentina: Museo Argentino de Ciencias Naturales (N.S.) Revista, v. 6, no. 2, p. 1–6. De Ferraris, C., 1947, Edad del arco o dorsal Antigua del Neuquén Oriental de acuerdo con la estratigrafía de la zona inmediata: Revista de la Sociedad Geológica Argentina, v. 2, no. 3, p. 256–283.
Digregorio, J.H., and Uliana, M.A., 1980, Cuenca Neuquina, in Turner, J.C.M., ed., Segundo Simposio de Geología Regional Argentina, Córdoba: Academia Nacional de Ciencias v. 2, p. 985–1032. Eisner, P., 1991, Tectonostratigraphic evolution of Neuquén Basin, Argentina [Master’s thesis]: Houston, Rice University, 56 p. Fernández, M.L., Verzi, H., and Sanchez, E., 2003, Actividad tectónica y evolución sedimentaria de los depósitos Thitoniano/Valanginianos tempranos, porción oriental de la Cuenca Neuquén, Argentina, in 8º Simposio Bolivariano–Exploración en Cuencas Subandinas Cartagena, p. 243–246 (CD-ROM). Franzese, J.R., and Spalletti, L.A., 2001, Late Triassic continental extension in southwestern Gondwana: Tectonic segmentation and pre-break-up rifting: Journal of South American Earth Sciences, v. 14, p. 257–270, doi: 10.1016/S0895-9811(01)00029-3. Gómez Omil, R., Schmithalter, J., Cangini, A., Albariño, L., and Corsi, A., 2002, El Grupo Cuyo en la dorsal de Huincul, consideraciones estratigráficas, tectónicas y petroleras, Cuenca Neuquina, in 5 Congreso de Exploración y Desarrollo de Hidrocarburos (Mar del Plata), Actas (CD-ROM).). González Díaz, E.F., and Nullo, F.E., 1980, Cordillera Neuquina, in Turner, J.C.M., ed., Segundo Simposio de Geología Regional Argentina, Córdoba: Academia Nacional de Ciencias, v. 2, p. 1099–1147. Groeber, P., 1929, Líneas fundamentales de la geología del Neuquén, sur de Mendoza y regiones adyacentes: Dirección Nacional de Geología y Minería, Publicación, v. 58, p. 1–110. Gulisano, C.A., Gutierrez Pleimling, A.R. and Digregorio, R.E., 1984, Esquema estratigráfico de la secuencia Jurásica al oeste de la provincia del Neuquén, in 9 Congreso Geológico Argentino, Buenos Aires: Actas, v. 1, p. 236–259. Harding, T.P., 1973, Areas potencialmente explotables para hidrocarburos resultantes de deformaciones producidas por fallas de desplazamiento lateral: Buenos Aires, Instituto Argentino del Petróleo, Resumen, 1 p. Harding, T.P., 1974, Acumulaciones importantes de hidrocarburos originadas por deformaciones causadas por fallas laterales: Petrotecnia, v. 2, no. 17, p. 12–19. Herrero Ducloux, A., 1946, Contribución al conocimiento geológico del Neuquén extrandino: Boletín Informaciones Petroleras, v. 23, no. 226, p. 245–281. Jaillard, E., Soler, P., Carlier, C., and Mourier, T., 1990, Geodynamic evolution of the northern and central Andes during the middle Mesozoic times; a Tethyan model: Journal of the Geological Society of London, v. 147, p. 1009–1022. 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. Kleiman, L.E., 2002, Magmatism and tectonic evolution of the Choiyoi and Puesto Viejo volcanics (Late Paleozoic–Early Mesozoic) at 34–35°S latitude, San Rafael, Mendoza, Argentina, in 15 Congreso Geológico Argentino (El Calafate): Actas, v. 2, p. 15–16. Mosquera, A., 2002, Inversión tectónica Jurásica inferior, sector central de la dorsal de Huincul, Área Los Bastos, in 5 Congreso de Exploración y Desarrollo de Hidrocarburos (Mar del Plata): Actas (CD-ROM). Orchuela, I.A., and Ploszkiewicz, J.V., 1984, La Cuenca Neuquina, in Ramos, V.A., ed., Geología y recursos naturales de la provincia de Río Negro: 9 Congreso Geológico Argentino (San Carlos de Bariloche), Relatorio, v. 1, no. 7, p. 163–188. Orchuela, I.A., Ploszkiewicz, J.V., and Viñes, R.F., 1981, Reinterpretación estructural de la denominada dorsal Neuquina, in 8 Congreso Geológico Argentino (San Luis): Actas, v. 3, p. 281–293. Pángaro, F. and Bruveris, P., 1999, Reactivación multiepisódica de sistemas extesionales, cuenca neuquina, Argentina: 14º Congreso Geológico Argentino, Actas 1, p. 231–234, Salta. Pángaro, F., Veiga, R., and Vergani, G., 2002, Evolución tecto-sedimentaria del area de Cerro Bandera, Cuenca Neuquina, Argentina, in 5 Congreso Argentino de Exploración y Desarrollo de Hidrocarburos (Mar del Plata): Actas (CD-ROM).
Intraplate deformation in the Neuquén Embayment Pardo Casas, F., and Molnar, P., 1987, Relative motion of the Nazca (Farallon) and South American plates since Late Cretaceous time: Tectonics, v. 6, no. 3, p. 233–248. Ploszkiewicz, J.V., Orchuela, I.A., Vaillard, J.C., and Viñes, R.F., 1984, Compresión y desplazamiento lateral en la zona de Falla Huincul: Estructuras asociadas, provincia del Neuquén, in 9 Congreso Geológico Argentino (San Carlos de Bariloche): Actas, v. 2, p. 163–169. Ramos, V.A., 1977, Estructura, in Rolleri, E.O., ed., Geología y recursos naturales del Neuquén: 7 Congreso Geológico Argentino, Relatorio, p. 99–118. Ramos, V.A., 1981, Descripción geológica de la hoja 33c Los Chihuidos Norte, Provincia del Neuquén: Servicio Geológico Nacional, Boletín, v. 182, p. 1–103. Ramos, V.A., 1984, Patagonia: ¿Un continente Paleozoico a la deriva?, in 9 Congreso Geológico Argentino: Actas, v. 2, p. 311–325. Ramos, V.A., 1996, Evolución tectónica de la Plataforma Continental, in Ramos, V.A., and Turic, M.A., eds., Geología y recursos naturales de la plataforma Continental Argentina: Buenos Aires, Asociación Geólogica Argentina e Instituto Argentino del Petróleo, p. 385–404. Ramos, V.A., 2000, Evolución tectónica de la Argentina, in Caminos, R., ed., Geología Argentina: Buenos Aires, Instituto de Geología y Recursos Minerales, Anales, v. 29, no. 24, p. 715–784. Ramos, V.A., 2004a, La plataforma Patagónica y sus relaciones con la plataforma Brasilera (chapter 22), in Mantesso-Neto, V., Bartorelli, A., Ré Carneiro, C.D., and Brito Neves, B.B., eds., Geologia do continente Sul-Americano: Sao Paulo, Beca Producoes Culturais Ltda., p. 371–381. Ramos, V.A., 2004b, Cuyania, an exotic block to Gondwana: Review of a historical success and the present problems: Gondwana Research, v. 7, no. 4, p. 1009–1026, doi: 10.1016/S1342-937X(05)71081-9. Ramos, V.A., and Barbieri, M., 1989, El volcanismo Cenozoico de Huantraico: Edad y relaciones isotópicas iniciales, provincia del Neuquén: Asociación Geológica Argentina, Revista, v. 53, no. 2, 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.D., Spalletti, L., Howell, J.A., and Schwarz, E., eds., The Neuquén Basin: A case study in sequence stratigraphy and basin dynamics: Geological Society of London Special Publication 252, p. 15–35. Ramos, V.A., and Kay, S.M., 2006, this volume, Overview of the tectonic evolution of the southern Central Andes of Mendoza and Neuquén (35°–39°S latitude), 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 latitude): Geological Society of America Special Paper 407, doi: 10.1130/2006.2407(01). 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, in 9 Congreso Geológico Argentino (Bariloche): Actas, v. 2, p. 84–106. Ramos, V.A., Wienecke, S., and Götze, H., 2002, El basamento de la cuenca Neuquina y regiones adyacentes: Datos gravimétricos preliminaries, in 5 Congreso de Exploración y Desarrollo de Hidrocarburos (Mar del Plata): Actas (CD-ROM). Scheuber, E., Bogdanic, T., Jensen, A., and Reutter, K.-J., 1994, Tectonic development of the North Chilean Andes in relation to plate convergence and magmatism since the Jurassic, in Reutter, K.-J., Scheuber, E., and Wigger, P.J., eds., Tectonics of the southern Central Andes: Structure and evolution of an active continental margin: Berlin, Springer-Verlag, p. 121–139. Somoza, R., 1998, Updated Nazca (Farallon)–South America relative motions during the last 40 my: Implications for mountain building in the central Andean region: Journal of South American Earth Sciences, v. 11, no. 3, p. 211–215, doi: 10.1016/S0895-9811(98)00012-1.
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MANUSCRIPT ACCEPTED BY THE SOCIETY 22 DECEMBER 2005
Printed in the USA
Geological Society of America Special Paper 407 2006
Structural evolution and magmatic characteristics of the Agrio fold-and-thrust belt Gonzalo Zamora Valcarce* Tomás Zapata* Repsol YPF, Faja Plegada Sur, Esmeralda 255, of. 1001, Buenos Aires C1035ABE, Argentina Daniel del Pino* Estudios Geológicos, Domene 414, Neuquén Q8300QBF, Argentina Andrés Ansa* Repsol YPF, Faja Plegada Sur, Esmeralda 255, of. 1001, Buenos Aires C1035ABE, Argentina
ABSTRACT The Agrio fold-and-thrust belt is located between 37°S and 38°S latitude in the eastern part of the Neuquén Andes. The belt can be divided into a western inner sector and an eastern outer sector. The inner sector is characterized by a thick-skinned deformation style. The dominant structures are large anticlines produced by the inversion of halfgrabens formed during the Triassic-Jurassic extension that initiated the Neuquén Basin. The outer sector is characterized by thin-skinned structures; recent studies have shown that these structures have been reactivated in a thick-skinned style. A long-standing question has been whether the deformation in this belt occurred in a continuous pulse or in two independent pulses. The analyses of synorogenic deposits, crosscutting relationships between magmatic rocks and sedimentary formations, and new single-crystal 40Ar/ 39Ar ages from volcanic rocks presented here indicate a minimum age of 102 Ma for the beginning of deformation in this belt and that deformation occurred in at least two pulses, one during the Lower to Middle Cretaceous, and a second one in the middle Miocene, with different degrees of propagation into the foreland. Keywords: Agrio fold-and-thrust belt, inversion, deformation stages, Ar-Ar ages, synorogenic deposits.
INTRODUCTION The age and temporal history of deformation in the Agrio fold-and-thrust belt have been topics of long discussion in the literature on the Neuquén Basin and the south-central Andes *E-mails:
[email protected];
[email protected]; delpinus@ infovia.com.ar;
[email protected].
(Ramos, 1978, 1998; Zapata et al., 2002, 2003; Repol et al., 2002; Cobbold and Rossello, 2003). The presence of Miocene synorogenic deposits has been argued to constrain the age of the youngest deformation (Ramos, 1998; Zapata et al., 2002, 2003), but the age of initial deformation and the number of distinct deformational pulses are still widely debated. The purpose of this paper is to present new single-crystal 40Ar/ 39Ar ages from volcanic rocks, field relationships between igneous and sedi-
Zamora Valcarce, G., Zapata, T., del Pino, D., and Ansa, A., 2006, Structural evolution and magmatic characteristics of the Agrio fold-and-thrust 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, doi: 10.1130/2006.2407(06). For permission to copy, contact
[email protected]. ©2006 Geological Society of America. All rights reserved.
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mentary rocks, data from seismic lines and exploration wells, and geochemical data on magmatic rocks that constrain the initial age of deformation and confirm the existence of two deformational pulses. GEOLOGIC AND TECTONIC SETTING The study area is located in the northwestern sector of Neuquén province in Argentina between 37° and 38°S latitude (Fig. 1). The region of interest lies in the western part of the Neuquén Basin and is part of the Neuquén fold-and-thrust belt in the Principal Cordillera. The Agrio fold-and-thrust belt, as defined by Bracaccini (1970), consists of a series of faultbounded basement blocks and large double-plunging anticlines developed in a Mesozoic sedimentary sequence. Based on morphostructural characteristics, Ramos (1978) divided the Agrio fold-and-thrust belt into an internal sector in the west and an external sector in the east. The internal sector is dominated by a thick-skinned deformational style, whereas the external sector is dominated by large anticlines associated with detachment folds. To the west, the Agrio belt is bounded by the extensionally active Loncopué trough, which is filled by a thick pile of Quaternary volcanic rocks. The eastern limit of the Agrio belt corresponds with the broad and gently dipping anticline of the Los Chihuidos high. To the north, the Agrio fold-and-thrust belt is separated from the Chos Malal trough and the Cordillera del Viento by the Cortaderas lineament. The Cortaderas lineament has a subtle expression on the surface but can be recognized as a prominent basement structure in seismic profiles. The Agrio fold-and-thrust belt lies in the Neuquén Basin, the history of which can be considered in three stages: (1) a Triassic to Early Jurassic prerift and rift stage, (2) a Late Jurassic to Cretaceous subsidence stage, and (3) a Tertiary to Holocene modification stage punctuated by magmatic events. Tectonic summaries (Kozlowski et al., 1993; Uliana and Legarreta, 1993; Manceda and Figueroa, 1995; Vergani et al., 1995) emphasize the first two stages. The Tertiary magmatic and tectonic evolution is the least understood and little discussed up to this point. This final stage of the evolution of the Neuquén Basin is a consequence of expansion and retreats of the magmatic arc (see summaries in Kay et al., this volume, chapter 2; Ramos and Kay, this volume, chapter 1; Ramos and Folguera, 2006). The Mesozoic and Paleogene sedimentary deposits of the Neuquén Basin constitute a sequence that can be up to 8000 m thick. These deposits are locally covered by Neogene synorogenic deposits. The Agrio fold-and-thrust belt occurs in the basin depocenter (Legarreta and Gulisano, 1989) where Jurassic to Upper Cretaceous sedimentary strata are well represented (Fig. 2). The basement of the Neuquén Basin crops out in the Cordillera del Viento, north of the study area (Zollner and Amos, 1973). It is composed of volcanic and volcaniclastic deposits of the Choiyoi Group of Permian-Triassic age, which are associated with crustal extensional on a regional scale (Ramos and Kay, 1991), and the sedimentary and volcanic
rocks of the Carboniferous Andacollo Group. The oldest basement outcrops in the study area are marine deposits of the Cuyo and Lotena Groups. They are found in the proximity of the town of Loncopué (Fig. 2; Zavala et al., 2002), and they have been documented in several oil exploration wells. The top of these marine sequences corresponds to the evaporites of the Auquilco Formation, which constitutes the basal décollement of the fold-and-thrust belt (Viñes, 1985). Overlying the basement are Kimmeridgian to Early Cretaceous sedimentary sequences (see Fig. 3). These sequences start with the marine deposits of the Mendoza Group, which conformably overlie the older rocks. The continental sandstones of the Tordillo Formation (Gulisano Gutierrez Pleimling, 1994) comprise the base of this group. Above them is the Vaca Muerta–Quintuco Formation (Fig. 3), which is a basinal deposit characterized by a finely stratified alternation of black and gray shales, calcareous micritic limestones, and bituminous marls. Because of its high content in organic matter, it is an excellent source of hydrocarbons and the most prolific source rock in the Neuquén Basin. The next unit is the Mulichinco Formation, which corresponds to a sandy platform deposit (Zavala, 2000). A new transgressive sequence, corresponding to the Agrio Formation, overlies it. The intercalated Avilé Member represents a sea-level lowstand sequence. The Huitrín Formation overlying this unit is composed of the eolian and river sandstones of the Lower Troncoso Member and evaporites of the Upper Troncoso Member. These evaporites, which are widely distributed across the basin, constitute the regional upper décollement of the Agrio fold-and-thrust belt (Ploszkiewicz, 1987). The evaporites mark the desiccation of the basin (Legarreta and Gulisano, 1989) that terminated with the deposits of the La Tosca Member, which record the last marine event in the basin. Subsequently, deposits of evaporites and clastic sedimentary rocks corresponding to the Rayoso Group were deposited as the basin shallowed. They represent the culmination of the Andean Cycle (Vergani et al., 1995). Recent studies of the Rayoso Group in the study area have documented a discordance within the group (Zavala et al., 2002). Continental red sandstones, conglomerates, and shales of the Neuquén Group cover these beds and mark the continentalization of the basin. They are exposed in the eastern part of the study area, where they begin with basaltic conglomerates levels and rhyolitic tuffs that may be derived from the Cordillera del Viento to the northwest. A series of Neogene deposits that fill small basins is well exposed in the study area. They have been recognized in two areas. The first is on the west side of Cerro Naunauco (Fig. 2), where Ramos (1998) described the Tralalhué Conglomerate, which was interpreted as part of a Miocene piggyback basin. This unit consists of a conglomeratic sequence composed of angular clasts from andesites, dacites, and calcareous subrounded clasts. Recent mammal fossils confirm its Miocene age (Repol et al., 2002). Field work in this study area shows that the Tralalhué Conglomerate unconformity overlies the Cretaceous Rayoso Group and the igneous rocks of Cerro Naunauco. The
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Figure 1. Regional location map of the Agrio fold-and-thrust belt, where the main morphostructural units of the Neuquén Basin are displayed.
second area is in the Pampa de Agua Amarga (Fig. 2), where the Puesto Burgos Formation occurs. This unit is composed of primary and reworked pyroclastic deposits and tuffs (Leanza and Hugo, 2001). These gently folded deposits unconformably overlie the Neuquén Group. On the basis of mammal fossils, Zapata et al. (2002) argued that these synorogenic deposits are of middle Miocene age. These deposits are onlapped by the Rincón Bayo Formation, which marks the end of deformation.
AGRIO FOLD-AND-THRUST BELT The Agrio fold-and-thrust belt, which was called the Agrio trough by Bracaccini (1970), is mainly characterized by detachment folds. The geological map of the study area (Fig. 2) shows N-NW–trending structures, with a difference in structural styles between the western and the eastern parts. Broad anticlines, products of basement inversion, characterize the structure to the
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Figure 2. Geological map of the study area.
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west (Ramos, 1998). To the east, the fold-and-thrust belt is characterized by large double-plunging anticlines separated by broad rhombic-shaped synclines, reflecting the presence of basement blocks at depth (Ramos, 1978; Viñes, 1985; Zapata et al., 1999a, 1999b, 2002). Traditionally, the Agrio fold-and-thrust belt has been divided into three regions (Ramos, 1998): the Loncopué graben, and the inner and outer sectors, structured by thick- and thin- skinned structure, respectively. Recent studies have demonstrated that the entire Agrio fold-and-thrust belt has undergone thick-skinned deformation (Zapata et al., 1999a, 1999b, 2002, 2003). This foldand-thrust belt is bounded to the north by the Cortaderas lineament, and to the east by the Chihuidos high (Fig. 4). The following analysis of the subsurface structure is based on surface geology derived from geological maps by YPF (Yacimientos Petroleros Fiscales) geologists and new observations, and on the interpretation of seismic sections (Figs. 7–9) that cover the transition between the inner and the outer sectors. They show excellent examples of the basement inversion and the Miocene synorogenic deposits. Loncopué Fault System The Loncopué trough is a 300-km-long and 30–40-kmwide depression that bounds the western side of the Agrio foldand-thrust belt. Ramos (1978) defined it as an extensional system. Numerous monogenic basaltic cones and Tertiary volcaniclastic and taphrogenic series fill it, but its initial fill is unknown. Because the thickest section of the volcanic rocks in the Eocene Cayanta Formation are found in this graben, it is assumed that the Loncopué trough existed in the Paleogene (Llambías and Rapela, 1989). Because geophysical data are lacking, there is no consensus on the geometry of the boundary between the Loncopué through and the Agrio fold-and-thrust belt. Eisner (1991), Lesta et al. (1985), Ramos (1998), and Cobbold and Rossello (2003) favor an eastward-dipping thrust, whereas Ramos (1978), Zapata et al. (1999a, 1999b), and Jordan et al. (2001) propose a west-dipping normal fault. Inner Sector
Figure 3. Schematic stratigraphic column (without scale) that shows the main lithologic units of the region, modified from Brissón and Veiga (1998).
The inner sector corresponds to the western part of the Agrio fold-and-thrust belt (Fig. 4). Vergani et al. (1995) interpreted the basement faults in this region as part of the Tres Chorros extensional system, which controlled the Jurassic depocenters of the Neuquén Basin. This zone likely correlates with the Cordillera del Viento (Zapata et al., 1999a, 1999b; Zapata and Folguera, 2006), which is related to the inversion of another Jurassic depocenter. The analysis of folding and associated faults, as well as their morphological expression in the area, has allowed the recognition of a series of basement highs and lows (Viñes, 1985; Ramos, 1998; Zapata et al., 1999a, 1999b, 2002) with N-NW axes (Fig. 5). These axes can be related to basement
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Figure 4. Regions of the Agrio fold-and-thrust belt and its limits to the east and to the north.
inversion, probably associated with older Jurassic depocenters (Vergani et al., 1995). The most representative of these structures in the study area is the Cerro Mocho anticline (Figs. 5 and 6). Even though there is no direct evidence for such a depocenter, the fact that the CMO X-1 exploration well, emplaced on the eastern side of the Cerro Mocho anticline (Fig. 6), penetrated more than 1500 m of black shales from the Los Molles Formation documents the presence of an anomalous thickness best interpreted as a Jurassic depocenter (Zapata et al., 2002). The oldest rocks of the region crop out near the core of the basement-related structures, as documented by the Cuyo Group rocks found on the western flank of the Cerro Mocho anticline (Fig. 2). The basement-related faults do not crop out; they can only be recognized on seismic sections. Figure 7 shows two seismic lines (time sections) located on the eastern boundary of the Cerro Mocho anticline. The first package of reflectors corresponds to the Mendoza and the Rayoso Group, drilled by the PDS X-1 exploration well, and serves as a primary input to tie the geology with the seismic information. Below these reflectors, a series of west-dipping inverted faults can be easily recognized. They are interpreted as the inversion of previous
Jurassic half-grabens. These faults do not cut upsection. Instead, they are inserted into the evaporite levels of the Auquilco Formation (documented by the PDS X-1 exploration well), transferring horizontal shortening to the frontal part of the Agrio fold-and-thrust belt. Outer Sector Interactions between adjacent anticlinal domes, such as those between the Pichi Mula, La Mula–Naunauco, and Chorriaca anticlines (Figs. 5, 6, and 8), are common in outer sector structures. Although there are few places to analyze whether a fault or a squeezed syncline separates the anticlines, field work could corroborate the existence of a fault between the adjacent anticlines for stratigraphic levels older than the evaporite of the Upper Troncoso Member. The seismic sections show a basal detachment near the top of the evaporite of the Auquilco Formation (Late Jurassic; Viñes, 1985; Figs. 7 and 8) and an upper detachment level in the Huitrín Formation (Early Cretaceous) (Fig. 8). The resulting geometry of the structures in this area is a typical fault-bend fold, with some internal detachment folds,
Structural evolution and magmatic characteristics of the Agrio fold-and-thrust belt
Figure 5. Landsat TM image of the Agrio fold-and-thrust belt showing the main fold axis and the morphostructural elements in the study area.
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Figure 6. East-west structural cross section through the Agrio fold-and-thrust belt (see location in Fig. 5). The structure of the Agrio fold-andthrust belt combines thin-skinned and thick-skinned deformation (modified from Zapata et al., 1999a, 1999b).
known in the area as the “Chorriaca fold type” (see Figs. 6 and 8; Groeber, 1946a; Herrero Ducloux, 1946). In the triangle zone shown on Figure 8, the Pichi Mula anticline could not evolve to a typical Chorriaca fold type, because it grew as a fold propagation fault through the eastern limb of the Cerro La Mula– Naunauco anticline, which was formed first. When the basement inversion of the frontal part of the Agrio fold-and-thrust belt occurred later, deep basement faults cut the previous folds (Fig. 8). At this time, some of the detachment folds evolved to either a fold-bend fault or a fold-propagation fault geometry, as seen in the Pichi Mula triangle zone (Fig. 8; Zapata et al., 2002). Such refolded triangle zones characterize the external structure of the Agrio fold-and-thrust belt. This inversion corresponds to a second deformation stage that produced the uplift of the triangle zone and the synorogenic deposits at the back limb of the Pichi Mula anticline (Puesto Burgos Formation). These deposits have been analyzed in the Pampa de Agua Amarga area (Fig. 2), where an onlap relationship can be observed between the Puesto Burgos and the sandstones and shales of the Neuquén Group (Fig. 9A). Figure 9B shows a seismic line (time) crossing the external part of the Agrio fold-and-thrust belt, where the same relationship can be observed. The age of these units constrain the timing of the deformation to the middle Miocene (Zapata et al., 2002; Zamora Valcarce et al., 2005). MAGMATIC ROCKS Little information is available on the chemistry and ages of Mesozoic igneous rocks at the latitude of the study area. Data from these rocks are important in establishing the age of uplift and the timing of deformation in the area. Volcanic and subvolcanic rocks in the study area occur in the inner part of the Agrio fold-and-thrust belt, where they are mainly found in the Collipilli volcanic field and the Cerro Naunauco laccolith (Figs. 2 and 10). In addition, a series of E-W–trending basaltic dikes cut the
Cerro Mocho anticline (Figs. 2 and 10). The Collipilli Group and the Cerro Mocho dikes were included in the same unit by Leanza and Hugo (2001), but the new single-crystal 40Ar/ 39Ar ages in Figure 11 show that the Cerro Mocho dikes have an age of ca. 100 Ma, and the Collipilli volcanic rocks are from 76 to 63 Ma in age. Crosscutting relationships of volcanic units with sedimentary rocks, along with the 40Ar/ 39Ar ages, constrain the deformation age of the Agrio fold-and-thrust belt. Collipilli Volcanic Area Magmatic rocks in the Collipilli area consist of volcanic facies associated with cones, lava flows, dikes, sills, and laccoliths. Llambías and Malvicini (1978) were the first to describe Collipilli area volcanic rocks. Later, Llambías and Rapela (1987) included them in the Neuquino-Mendocina volcanic province, which encompasses units between 38°30′S and 34°S (Groeber, 1946a, 1946b; Yrigoyen, 1972; Bettini, 1982; Kozlowski et al., 1987; Haller et al., 1985). Domes and laccoliths characterize these series. The intrusive series of the Collipilli Formation were emplaced as laccoliths and associated sills. These laccoliths intruded the contact between the Agrio and the Rayoso Formations, filling the space of 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 related to extrusive domes, pyroclastic flow deposits, and massive lava flows. These volcanic rocks unconformably cover the Agrio, Huitrín, and Rayoso Formations, which are folded in the Collipilli syncline (Figs. 2 and 10). Therefore, the area was already uplifted (and partially eroded) before or at least simultaneously with their intrusion. In several places, there are conglomeratic sandstones and volcanogenic deposits intercalated between the subvolcanic volcanic rocks. There is also evidence of younger deformation, as indicated by small thrust faults that cut these
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Figure 7. Seismic sections showing the basement faults being inserted into the Auquilco evaporites and transferring the shortening to the outer sector. See location in Figure 5. TWT—two-way traveltime.
rocks. These thrusts may be a consequence of the Miocene regional uplift discussed next. 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, for which known ages ranged from 50 to 45 Ma (Llambías and Rapela, 1989), and (2) an Eocene volcanic facies called the Cayanta Formation (Rapela and Llambías, 1985), which had one K/Ar age of 39 ± 9.11 Ma (Llambías and Rapela, 1989). They noted that volcanic rocks mapped in the
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Figure 8. Detail of east-west cross section of the outer zone of the Agrio fold-and-thrust belt showing the two detachments levels and basement reactivation fault (modified from Zapata et al. 2002). See location and legend in Figure 6.
Molle Formation farther north yielded K-Ar whole-rock ages of 71.5 ± 5 Ma (Llambías et al., 1978), so that not all volcanic rocks in the Molle Formation could be Eocene in age. More problems for assigning an Eocene age to Molle Formation volcanic centers arise from recently reported ages, such as the Miocene 40Ar/ 39Ar age of 11.7 ± 0.2 Ma for the Cerro Negro center, west of Chos Malal (Kay, 2001; Kay et al., this volume, chapter 2), the 40Ar/ 39Ar age of 60.1 ± 1.6 Ma from igneous rocks of Cerro Nevazón (eastern flank of the Cordillera del Viento; Franchini et al., 2003), and the K/Ar ages of 60.7 ± 1.9 Ma in Campana Mahuida (south of the study area; Franchini et al., 2003). Sillitoe (1977) reported ages of 74.2 ± 1.4 Ma for flat-lying andesitic porphyry intruded into previously folded Jurassic sediment in Campana Mahuida. New 40Ar/ 39Ar single-crystal analyses (Table 1) yield welldefined plateaus (Figs. 11A and 11B) that indicate ages of 65.5 ± 0.46 Ma for the Cerro Naunauco laccolith, 72.83 ± 0.83 Ma for a volcanic bomb from Collipilli region, and 56.64 ± 0.44 Ma for an andesitic sill (Fig. 11C) emplaced in the Agrio Formation in the Collipilli region. The first two ages overlap the 65–75 Ma age range that has been reported for andesitic volcanic rocks uncomformably overlying the Tordillo Formation on the western flank of the Cordillera del Viento anticline (Linares and Gonzalez, 1990; Franchini and Schalamuk, 1999). Geochemistry Chemical analyses of the Collipilli and Naunauco samples (Table 2) show that they are subalkaline low-K andesites with Nb/Y ratios that vary between 0.2 and 0.3 (Fig. 12). The analyses
fall in the volcanic arc field, which is typical of the Andean margin, on trace-element discriminative diagrams; this is represented by their arc- to backarc-like signature (La/Ta = 66; Ba/La = 21; Ta/Hf = 0.10; Fig. 13). Other characteristics include relatively flat rare earth element (REE) patterns (La/Yb = 10; La/Sm = 5.5; Sm/Yb = 2.2). Even though the rocks of Collipilli and the Cerro Naunauco have different degrees of differentiation, the trace elements and incompatible rare earth elements indicate a common source (Table 1; Fig. 13). The Collipilli and Naunauco volcanic rocks analyzed here are geochemically similar to those analyzed by Llambías and Malvicini (1978). The main difference is that the new Ar-Ar ages indicate that these Collipilli igneous rocks need to be reassigned to an older igneous event. The Collipilli Group is proposed here to designate these rocks. The differences in the Ar/Ar ages presented here and the previous K/Ar ages could indicate two different events, but they also might be interpreted as a long-lasting igneous cycle that extended from ca. 70 Ma until 39 ± 9 Ma (age from Cerro Mayal; Cobbold and Rossello, 2003). More data are needed to define the Eocene magmatic event in this area of the Andes. Dikes of Cerro Mocho Leanza and Hugo (2001) mapped the dikes in the Cerro Mocho region as a subunit of the Collipilli Formation. These dikes are exposed as a series of E-W dikes emplaced along preexisting structures. The most distinctive is a 19-km-long dike that cuts E-W through the Cerro Mocho anticline. It is better seen on
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Figure 9. (A) View to the north of a field example of the synorogenic deposits of the Puesto Burgos Formation; note the onlap relationship with the sediments of the Neuquén Group. See location in Figure 8. (B) Seismic section (time) in the Pampa de Agua Amarga showing the unconformity of the synorogenic deposits over the Neuquén Group sediments. See location in Figure 5. TWT—two-way traveltime.
the Landsat image (Figs. 5 and 10) than in the field due its lack of topographic expression. In detail, it is not a single dike, but a series of dike segments (Fig. 10) with chilled margins. Repol et al. (2002) put the Cerro Mocho dikes into 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 Rosello, 2001). Two sam-
ples selected for single-crystal 40Ar/ 39Ar ages (Table 1) yielded the plateaus shown in Figure 11D. The first, sample M3-9 yielded a well-defined plateau indicating an age of 101.99 ± 0.69 Ma (late Early Cretaceous–Albian). A second sample does not show a well-defined plateau because of its alteration, but an age older than those obtained from the Collipilli and Naunauco igneous rocks can be inferred (around 90 ± 4.06 Ma), and an
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Figure 10. (A) Outcrop in the Collipilli area showing the western limb of the Collipilli syncline. Note that the Collipilli volcanic rocks are restricted to the core of the syncline. (B) Field example showing the Cerro Mocho dikes cutting the Agrio Formation.
alteration event at 51 Ma might have affected the sample. These ages can be interpreted as an older magmatic event in the area, near 90–100 Ma, not reported before. Geochemistry The Cerro Mocho dikes are geochemically distinct from the Collipilli and Naunauco samples because they are subalkaline basalts (Table 3) with relatively higher Nb/Y ratios
(0.3–0.4; Fig. 12). They have an intermediate signature (La/Ta = 49; Ba/La = 223; Ta/Hf = 0.10; Fig. 13) and a relatively flat REE pattern (La/Yb = 12.2; La/Sm = 2.26; Sm/Yb = 5.39) showing mantle affinity. The dikes of the Cerro Mocho area are distinctive, even though they maintain characteristics associated with a convergent continental margin. They have a minor Nb anomaly with respect to Th and Ce; the high Ti and Y contents also outline a transitional
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Figure 11. Age spectra and isochron plots for the 40Ar/ 39Ar single-crystal analysis from the samples in this study: (A) andesitic volcanic rock from Naunauco, (B) andesitic sill, (C) volcanic bomb from Collipilli, and (D) two basaltic dikes from the Cerro Mocho. MSWD—mean square of weighted deviates.
Step
T (C)
t (min.)
36
Ar
37
Ar
38
Ar
TABLE 1. DETAILS OF AR/AR DATING 40 Ar Ar %40Ar %39Ar rlsd
39
Ca/k
40
Ar*/39ArK
Age (Ma)
1 s.d.
M11-6 Naunauco Amphibole 16.68 mg (Preferred age: 65 Ma.) 1 850 12 25.488 18.730 5.011 4.224 2 950 12 1.412 14.620 0.360 3.555 3 990 12 0.627 16.010 0.160 2.456 4 1020 12 0.336 12.020 0.086 1.930 5 1050 12 0.404 33.340 0.164 5.364 6 1070 12 0.320 36.300 0.153 5.825 7 1095 12 1.202 152.260 0.591 24.529 8 1110 12 1.429 372.700 1.066 59.037 9 1125 12 0.384 179.920 0.454 28.837 10 1140 12 0.209 31.090 0.119 4.940 11 1160 12 0.295 31.270 0.135 4.738 12 1190 12 0.437 46.680 0.186 7.106 13 1220 12 0.426 55.340 0.189 8.397 14 1250 12 0.492 74.900 0.247 11.508 15 1280 12 0.268 40.640 0.140 6.181 16 1400 12 0.190 28.720 0.097 4.368
8003.570 551.080 253.850 151.850 241.450 227.660 897.680 1670.230 721.200 167.110 188.190 277.530 305.590 389.320 212.740 147.580
7.9 2.31 26.3 1.94 29.6 1.34 37.6 1.05 53.9 2.93 62.1 3.18 63.6 13.40 78.2 32.26 88.0 15.76 67.1 2.70 57.4 2.59 57.0 3.88 62.4 4.59 66.3 6.29 66.7 3.38 67.9 2.39 Cumulative %39AR rlsd = 100
16.04 14.87 23.63 22.58 22.53 22.59 22.50 22.88 22.61 22.81 23.93 23.82 23.90 23.60 23.84 23.84
151.99 405.681 40.95 118.576 30.55 89.183 29.24 85.455 24.21 71.034 24.18 70.952 23.45 68.843 22.31 65.571 22.15 65.095 22.41 65.857 22.57 66.317 22.20 65.236 22.68 66.629 22.44 65.953 22.79 66.950 22.37 65.735 Total gas age = 75.81 plateau age = 65.5 isochron age = 64.97
13.723 1.660 1.277 0.999 0.557 0.802 0.502 0.434 0.442 0.551 0.583 0.502 0.483 0.506 0.486 0.810 0.40 0.46 0.35
M1c-7 Sill - Plagioclase 23.60 mg (Preferred age: 56 Ma.) 1 650 12 32.638 53.145 7.489 99.911 2 730 12 11.122 87.900 3.166 85.568 3 810 12 6.466 78.649 2.298 86.356 4 890 12 17.360 80.506 5.220 151.402 5 960 12 10.449 56.497 3.962 153.992 6 1030 12 12.662 56.709 4.581 168.869 7 1100 12 17.085 73.774 5.219 159.366 8 1170 12 31.637 133.617 8.665 205.920 9 1220 12 9.198 66.895 2.245 36.931 10 1400 12 7.000 91.927 1.703 29.290
11652.000 4854.920 3521.500 7944.380 5960.480 6859.190 7940.450 13174.100 3430.480 2644.470
19.1 8.50 34.0 7.30 47.2 7.30 36.9 12.90 49.4 13.10 46.7 14.30 37.9 13.50 30.7 17.50 22.7 3.10 23.9 2.50 Cumulative %39AR rlsd = 100
1.76 3.41 3.02 1.76 1.22 1.11 1.53 2.15 6.01 10.43
22.35 65.660 19.39 57.080 19.35 56.960 19.49 57.370 19.23 56.630 19.08 56.190 18.98 55.910 19.74 58.090 21.20 62.320 21.73 63.860 Total gas age = 58 Plateau age = 56.64 steps (2-7) Isochron age = 56.1 steps (2-5)
1.050 0.580 0.500 0.550 0.460 0.460 0.540 0.630 0.910 0.860 0.42 0.44
43.93 126.54 25.7 75.09 24.33 71.17 24.55 71.81 24.67 72.14 25.3 73.95 27.55 80.40 34.77 100.88 39.39 113.86 35.51 102.98 Total gas age = 86.4 no plateau no isochron
2.18 0.84 0.65 0.78 0.87 1.03 1.10 2.74 3.59 2.07 0.70
M5-7 Collipilli Plagioclase 18.64 mg (Preferred age: 73 Ma.) 1 650 12 14.916 171.963 20.627 20.627 2 730 12 3.839 292.029 23.551 23.551 3 810 12 2.601 225.260 22.342 22.342 4 890 12 2.529 165.102 15.946 15.946 5 960 12 2.013 83.690 9.295 9.265 6 1030 12 1.344 53.448 7.397 7.397 7 1100 12 2.619 67.236 8.811 8.811 8 1160 12 5.026 67.513 6.055 6.055 9 1250 12 4.143 55.679 3.855 3.855 10 1400 12 4.008 115.353 5.760 5.760
5186.130 1669.150 1260.040 1096.760 798.176 567.226 989.042 1653.120 1341.020 1343.640
17.3 16.7 35.7 19.0 42.6 18.1 35.3 12.9 28.4 7.5 32.8 6.0 24.3 7.1 12.6 4.9 11.1 3.1 14.9 4.7 Cumulative %39AR rlsd = 100
30.08 44.94 36.45 37.44 32.51 26.04 27.51 40.36 52.47 73.19
0.95
M1-610 Cerro Mocho Ground mass 11.87 mg (Preferred age: 91.97 ± 4.06 Ma; Alternation between 57 and 43 Ma) 1 675 12 310.622 0.628 58.418 1.558 90902.670 0.5 6.100 2.3240083 304.6618 738.74 143.51 2 750 12 75.839 38.596 14.592 4.976 22166.995 0.4 19.600 45.298312 19.3867 57.18 16.03 3 820 12 24.768 60.568 4.820 3.539 7256.270 0.7 13.900 101.6207 14.6626 43.41 10.88 4 900 12 5.180 5.661 1.061 1.575 1543.660 2.3 6.200 20.838601 23.2033 68.22 4.33 5 980 12 17.705 4.905 3.711 3.826 5256.920 1.9 15.100 7.4029037 23.5108 69.11 3.6 6 1050 12 5.064 3.993 1.070 1.574 1510.740 2.4 6.200 14.68082 26.8425 78.69 4.94 7 1110 12 11.341 7.826 2.760 3.049 3423.420 3.6 12.000 14.854593 40.9922 118.84 4.44 8 1170 12 6.066 4.852 2.704 4.469 1883.440 6.3 17.600 6.2671546 26.7411 78.4 2.79 9 1400 12 1.143 0.675 0.522 0.831 351.124 4.0 3.300 4.6865857 18.376 53.26 4.12 Total Gas Age = 111.88 3.08 39 No Plateau Step 4 a 8 82.65 ± 4.02 Ma lost Ar = 57% 39 no isochron Step 6 a 8 91.97 ± 4.06 Ma lost Ar = 35.8% M3-9 Cerro Mocho Plagioclase 10.75 mg (Preferred age: 102 Ma) 1 570 12 7.974 11.078 1.662 7.462 2756.210 16.4 4.7 4.95 60.9 173.25 3.15 2 650 12 4.837 110.730 1.120 15.707 2020.500 31.4 9.8 23.63 40.85 118.03 1.26 3 730 12 2.381 84.985 0.689 19.465 1404.550 51.8 12.1 14.59 37.61 108.95 0.85 4 810 12 0.967 3.772 0.369 15.715 831.976 66.7 9.8 0.8 35.34 102.56 0.71 5 880 12 1.096 32.201 0.382 13.752 801.252 61.1 8.6 7.81 35.72 103.63 0.80 6 945 12 1.072 3.000 0.405 16.864 892.010 65.5 10.5 0.59 34.7 100.73 0.71 7 1010 12 1.173 3.046 0.394 13.429 806.810 58.2 8.4 0.76 35.01 101.62 0.78 8 1075 12 1.614 5.250 0.441 9.407 792.131 41.3 5.9 1.86 34.82 101.08 0.94 9 1140 12 3.078 10.338 0.726 10.508 1252.940 29.2 6.6 3.28 34.95 101.46 1.16 10 1195 12 6.182 15.421 1.548 24.995 2652.960 32.7 15.6 2.05 34.87 101.21 1.09 11 1260 12 3.890 9.715 0.936 11.099 1524.610 26.3 6.9 2.92 36.25 105.10 1.36 12 1400 12 0.642 2.077 0.159 1.986 258.113 28.9 1.2 3.48 37.09 107.49 1.26 Cumulative %39AR rlsd = 100 Total Gas Age = 107.83 Plateau age = 101.99 (steps 4-11) no isochron Note: Analyses were done at Actlabs. Analytical details can be found at http://www.actlabs.com/home.htm.
3.03 1.11 0.65 0.49 0.61 0.50 0.59 0.79 1.04 0.97 1.26 1.14
TABLE 2. MAJOR- AND TRACE-ELEMENT ANALYSES OF VOLCANIC ROCKS
Sample M9-6 M10-6 M11-6 M1B-7 M7-6 M5-6b M1-9 M5-8 M4-7 M5-7 M3-8 Major elements in wt% SiO2 51.5 52.29 49.08 53.61 52.88 51.34 51.02 48.81 48.73 47.25 51.19 51.76 56.09 58.44 TiO2 0.71 0.69 0.71 0.65 0.66 0.73 0.67 0.91 0.92 2.42 0.62 0.75 0.65 0.4 Al2O3 18.47 19.14 17.9 18.28 18.06 18.06 18.68 18.62 15.42 16.84 16.15 19.13 17.85 17.79 7.01 7.15 6.67 7.14 7.42 8.88 7.42 9.23 9.71 9.71 7.21 7.76 7.13 6.39 Fe2O3 MnO 0.08 0.05 0.06 0.15 0.17 0.18 0.06 0.13 0.15 0.08 0.14 0.17 0.16 0.15 MgO 4.46 3.9 4.29 3.5 3.23 3.78 3.82 5.67 8.67 5.26 3.08 2.34 2.12 2.65 CaO 7.88 8.01 8.21 3.8 5 5.33 7.6 10 9.73 5.06 7.38 8.66 8.87 6 Na2O 5.75 5.78 5.22 6.22 6.13 5.34 6.09 3.52 2.5 5.4 2.9 3.73 3.41 3.37 0.24 0.2 0.41 1.31 1.5 1.72 0.32 0.67 0.92 0.45 1.14 1.45 1.59 1.38 K2O P2O5 0.16 0.17 0.14 0.24 0.23 0.25 0.2 0.14 0.33 0.56 0.18 0.3 0.25 0.21 0.007 0.005 0.005 <.001 <.001 0.001 0.004 0.009 0.064 0.012 0.003 <.001 <.001 <.001 Cr2O3 LOI 3.1 2.7 7.6 4.8 4.6 4.1 3.8 2.3 2.8 6.9 9.7 3.7 1.6 3.1 TOT/C 0.03 0.04 0.08 0.33 0.5 0.03 0.02 < .01 0.01 0.86 2.55 0.02 0.27 0.03 TOT/S 0.01 0.01 0.01 0.01 0.01 0.06 0.05 0.02 0.01 0.09 0.11 0.01 0.01 < .01 Total 99.4 100.11 100.32 99.76 99.94 99.78 99.72 100.05 100 100.05 99.74 99.8 99.78 99.96 Trace elements in ppm Ga 20.3 20.4 19.8 21.2 20.1 23.2 22.4 19.6 14.9 20.5 18.3 23.2 21.8 18.7 Cs 1.3 0.9 1.2 0.4 0.6 0.6 0.7 1.1 1.3 0.3 2.1 1.4 1.8 1.6 Rb 4.2 2.4 8.3 40 42.8 44.6 9.4 13.8 24 13.4 35.7 36.6 48.6 40.2 Sr 668 716 610 852 821 643 604 479 504 446 484 904 674 616 Ba 151 187 203 419 458 520 230 242 237 884 390 438 494 568 Ta 0.2 0.3 0.2 0.3 0.3 0.3 0.2 0.2 0.4 0.3 0.2 0.3 0.4 0.3 Nb 3.3 3.5 3.4 4.5 4.6 4.9 3.5 2.9 4 6.8 3.3 5.4 5.9 4.3 Sc 23 20 22 15 15 18 20 33 31 21 15 15 14 9 Hf 2.2 2.1 2.1 3.3 3.8 2.7 2.3 1.9 1.9 3.1 2.1 3.9 3.6 2.9 Th 4.1 2.8 3.1 8.3 8.2 3.9 3.6 3 1.4 0.3 4 7.7 7.3 7.8 Zr 75.9 69.5 66 121.2 125.5 100.7 69.8 54.5 71.8 95.3 72.8 142.5 120.6 95.5 Y 16.9 14.8 15.4 17.5 17.5 18.8 16.4 16 20.4 16 13.7 24.7 18.9 13.7 Ni 25.7 17.7 20.3 14.3 10.3 47.3 < .1 27.9 50 62 8 6.9 10 87 Co 19.7 14.9 14.6 15 13.5 21 16.7 28.4 38 26.7 16.7 15 13.3 12.4 V 189 171 189 145 150 181 190 245 228 288 139 125 108 78 La 11.4 8.1 10 24.6 23.8 16.1 12.2 8.2 10.6 17.1 13.9 26 22.9 22.9 Ce 26.4 20.6 25.7 58.8 58.2 36.4 27.2 19.5 25.5 47.4 29.8 59.6 53.2 43.6 Pr 3.16 2.78 3.01 6.62 6.56 4.52 3.34 2.71 3.44 7.02 3.67 7.38 5.92 5.02 Nd 14.9 12.4 13.8 24.8 26.1 20.3 14.2 13.2 16.3 32.8 17.5 31.4 24.1 19.9 Sm 3.2 3.1 3 4.9 5.4 4.2 3.3 3.1 3.7 5.4 3 6.2 4.3 3.3 Eu 0.96 0.92 0.93 1.35 1.51 1.09 0.91 0.97 1.27 1.38 0.98 1.82 1.33 0.97 Gd 3.35 2.63 2.86 3.75 3.79 3.69 3.01 3.27 3.71 4.31 2.67 4.84 3.76 2.37 Tb 0.51 0.45 0.44 0.64 0.55 0.54 0.49 0.49 0.61 0.54 0.42 0.71 0.58 0.4 Dy 2.84 2.73 2.61 3.64 3.3 3.36 2.89 2.92 3.47 2.69 2.25 4.08 3.16 2.2 Ho 0.56 0.5 0.52 0.59 0.62 0.59 0.52 0.53 0.7 0.52 0.5 0.74 0.65 0.44 Er 1.48 1.61 1.61 1.79 1.95 1.75 1.5 1.54 1.83 1.25 1.21 2.38 1.71 1.27 Tm 0.24 0.24 0.23 0.26 0.29 0.29 0.23 0.24 0.26 0.19 0.19 0.41 0.29 0.2 Yb 1.72 1.6 1.57 1.71 1.95 1.96 1.49 1.68 1.84 1.08 1.44 2.46 1.94 1.49 Lu 0.27 0.24 0.22 0.28 0.26 0.29 0.24 0.24 0.3 0.16 0.2 0.35 0.3 0.24 Ni 45 34 36 21 28 60 68 36 163 68 <20 <20 29 65 Note: Analyses were done at Actlabs. Analytical details can be found at http://www.actlabs.com/home.htm. LOI—loss on ignition, TOT—total.
Naunauco M1C-7 M2-6 M5-6
Collipilli M4-8 51.07 0.76 17.98 8.2 0.13 3.77 7.08 5.36 1.24 0.25 0.002 4.2 0.09 < .01 100.09 20 1.4 31.2 531 405 0.2 4.5 20 3.4 7.5 124.1 22.1 2.3 19.2 178 23.4 52.5 6.21 27 5.5 1.38 4.42 0.68 3.5 0.7 2.22 0.32 2.13 0.36 <20
M2-8 57.87 0.61 17.05 6.43 0.12 2.16 7.32 3.29 1.61 0.23 0.003 2.9 0.05 < .01 99.65 22.3 2.4 47.8 998 523 0.4 7.1 12 3.9 7.4 128.4 20.4 2.1 10.5 121 26.7 56.1 6.42 26 5.2 1.37 4.1 0.62 3.33 0.61 2.03 0.33 2.27 0.32 <20
21.7 8.8 22.3 512 267 0.3 5 14 3.5 7.2 119.1 18.9 2.9 13.3 144 19.9 45.1 5.75 24 4.9 1.37 4.03 0.58 3.38 0.58 1.87 0.29 1.85 0.3 <20
51.23 0.66 18.12 7.04 0.18 3.03 7.65 3.48 0.69 0.23 0.002 7.2 0.08 0.01 99.54
M7-8
18.3 3.1 78.3 506 1639 0.4 5.7 18 3 10.5 112.1 16.6 11 15 158 21.7 47.6 5.13 18.5 3.8 1.03 3.36 0.51 2.84 0.53 1.67 0.25 1.7 0.26 39
55.81 0.67 17.7 7 0.19 2.95 6.05 4.03 2.65 0.21 0.003 2.3 0.17 0.01 99.75
M3-7
20.9 3.8 48.2 536 645 0.3 5.5 21 3.6 6 111.4 20.7 15 17.3 189 20 46 5.44 22 4.6 1.34 4.17 0.67 3.85 0.64 2.01 0.3 2.16 0.33 45
52.36 0.77 17.84 7.49 0.18 3.34 7.72 3.74 1.53 0.18 0.002 5.1 0.71 0.04 100.33
M3B-7
15.8 0.5 53.2 607 501 0.7 11.5 7 4.4 4.4 155.3 21.1 46 9.9 73 24.3 49.4 5.81 24.2 4.5 1.39 3.86 0.61 3.45 0.64 1.79 0.29 2.14 0.3 <20
61.94 0.58 16.54 5.36 0.13 2 4.73 4.09 2.01 0.23 <.001 2.4 0.1 0.02 100.07
M4-9
Structural evolution and magmatic characteristics of the Agrio fold-and-thrust belt 139
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Figure 12. Geochemical classification of Winchester and Floyd (modified by Pearce, 1996).
pattern between arc basalts and mid-ocean-ridge basalts (MORBs). The high content of heavy rare earth and incompatible elements (Table 3) indicates a different source than the Naunauco and Collipilli volcanic rocks. It seems that the Cerro Mocho volcanic rocks present clear differences in geochemical and chronological analyses and cannot be considered as part of the Collipilli Group. Tectonic Setting and Correlations of Magmatic Rocks Figure 14 shows a schematic map with the locations of the main igneous rocks for the Upper Cretaceous through Paleogene. The 40Ar/ 39Ar ages (65.5 ± 0.46; Fig. 11A) of igneous rocks from Collipilli and the ages of the younger sills (56.64 ± 0.44; Fig. 11) are similar to the radiometric ages obtained by others in the surrounding areas. North of the study area, Franchini et al. (2003) reported K/Ar ages of 56.0 ± 1.7 and 40Ar/ 39Ar ages of 60.1 ± 1.6 Ma for igneous rocks from Cerro Nevazón, and K/Ar ages of 60.7 ± 1.9 Ma in Campana Mahuida. Domínguez et al. (1984) reported a K/Ar whole-rock age of 67 ± 3 Ma for Los Maitenes–El Salvaje tonalite stock, and J.I.C.A./M.M.A.J. (2000) gave a K/Ar whole-rock age of 64.7 ± 3.2 Ma for the Varvarcó tonalite stock. These Paleocene ages are older than those obtained for a Collipilli intrusive (49.9 ± 3.2 and 48 Ma, whole-rock K/Ar age; Llambías and Rapela, 1989) and the microdiorite stock of Cerro Caicayén (44.7 ± 2.2 Ma, wholerock K/Ar age; Llambías and Rapela, 1989). However, they are younger than those obtained in samples from the Cerro Naunauco (72.83 ± 0.83 Ma; Table 1; Fig. 11C) and the andesite dikes of the Campana Mahuida district according to the hydrothermal biotite age (74.2 ± 1.4 Ma; Sillitoe, 1977). Nevertheless, all these radiometric ages are younger than those obtained during this study from the Cerro Mocho dikes (101.99 ± 0.69 and 91.97 ± 4.06 Ma; Table 1; Fig. 11D). The alteration event seen in one of the samples at an age of 51 Ma (Fig. 11D) corre-
lates with the middle Eocene radiometric ages of Cerro Caicayén (Llambías and Rapela, 1989) and the Collipilli sills sampled in this study (56.64 ± 0.44 Ma; Fig. 11B). Based on the available ages, Franchini et al. (2003) grouped the igneous rocks into three different magmatic cycles: Late Cretaceous (Campana Mahuida), Paleocene (Campana Mahuida and Cerro Nevazón), and middle Eocene (Caicayén and Collipilli). With the new data presented in this study, the Collipilli igneous rocks should be assigned to the Maastrichtian-Danian (CretaceousPaleocene boundary) magmatic event. There was magmatic activity during the middle Eocene at the Collipilli area, as indicated by the sills and the alteration event registered on one of the samples of the Cerro Mocho dikes. In addition, an older Middle Cretaceous magmatic event is recorded by the Cerro Mocho dikes. This event has not been recorded before for this part of the Andes, and because very few samples have been dated, new radiometric ages are needed to establish a detailed chronology for this event. According to field relationships, the Collipilli and Cerro Mocho magmas (Figs. 2 and 10) intruded the already formed structures of the Agrio fold-and-thrust belt. In addition, the igneous rocks from the Collipilli and Caicayén volcanic units have been interpreted as having been intruded close to the end of the main deformational event (Llambías and Malvicini, 1978; Minniti et al., 1986; Llambías and Rapela, 1989; Franchini, 1992). This consideration together with field observations and the ages presented here show that the Agrio fold-and-thrust belt was already deformed by the end of the Cretaceous. STRUCTURE AND MAGMATISM In order to develop a comprehensive understanding of the changes in the structural style and time of deformation in this sector of the Neuquén Andes, it is necessary to combine all of the available information.
Structural evolution and magmatic characteristics of the Agrio fold-and-thrust belt
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Figure 13. (A) Ba/Ta versus La/Ta ratios for the samples analyzed. (B) Th/Hf versus Ta/Hf ratios for the samples analyzed. La/Ta, Ba/La, Ba/Ta, and Th/Hf ratios can be used as relative measures of the importance of a source component associated with a subducting slab. Relative fields for values commonly observed in intraplate, backarc, and arc magmas are indicated. MORB—mid-ocean-ridge basalt.
Lower to Middle Cretaceous (First Compressional Stage) According to Ramos (1981), most of the inversion of the half-graben systems or basement blocks took place during the Late Cretaceous (Cenomanian). Zapata et al. (2002) confirmed this interpretation based on field observations. Cobbold and Rossello (2003) showed evidence for a Late Cretaceous deforma-
tion, something that the first geologists working in the basin (e.g., Groeber, 1929; Keidel, 1925; Wichmann, 1934) already postulated. North of the study area, Manceda and Figueroa (1995) inferred ages of 94 Ma for tectonic inversion on the Chilean side from the subsidence curve. The new 101 Ma Ar-Ar age from the Cerro Mocho dikes in this study further confirms a Middle Cretaceous deformational
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G. Zamora Valcarce et al. TABLE 3. MAJOR- AND TRACE-ELEMENT ANALYSES OF CERRO MOCHO DIKES Sample M2-9 M3-9 M1-9 M1-6 Major elements in wt% SiO2 46.18 47.69 47.25 39.72 2.44 2.36 2.42 2.95 TiO2 16.65 14.96 16.84 15.95 Al2O3 Fe2O3 11.25 9.91 9.71 11.52 MnO 0.17 0.16 0.08 0.06 MgO 5.92 5.97 5.26 6.64 CaO 7.13 5.38 5.06 7.86 4.47 4.14 5.4 4.09 Na2O 0.82 0.26 0.45 0.17 K2O 0.47 0.43 0.56 0.43 P2O5 Cr2O3 0.013 0.016 0.012 0.036 LOI 3.2 7.9 6.9 10.3 TOT/C 1.24 1.02 0.86 1.77 TOT/S 0.18 0.21 0.09 0.38 Total 99.44 99.7 100.05 99.79 Trace elements in ppm Ga 19.5 19.4 20.5 18.4 Cs 1.7 1.8 0.3 1.2 Rb 20.9 7.4 13.4 5.4 Sr 1866 867 446 129 Ba 6402 4544 884 354 Ta 0.3 0.3 0.3 0.2 Nb 6 5.3 6.8 5 Sc 28 28 21 34 Hf 2.8 2.2 3.1 2.8 Th 1.7 0.8 0.3 0.4 Zr 81.7 73.3 95.3 74.7 Y 16.9 16 16 16.3 Ni 70 83 68 169 Co 32.8 31.9 26.7 40.3 La 14.2 12.8 17.1 10.1 Ce 43.8 39.9 47.4 32.5 Pr 6.67 6.15 7.02 5.35 Nd 32 30.7 32.8 27.9 Sm 6.6 6.5 5.4 5.7 Eu 1.65 2.05 1.38 1.8 Gd 4.96 4.95 4.31 4.88 Tb 0.66 0.67 0.54 0.67 Dy 3.35 3.63 2.69 3.45 Ho 0.55 0.59 0.52 0.58 Er 1.54 1.47 1.25 1.27 Tm 0.23 0.18 0.19 0.2 Yb 1.22 1.06 1.08 1.13 Lu 0.19 0.19 0.16 0.17 Ni 59.7 63.1 63.5 175.3 Note: Analyses were done at Actlabs. Analytical details can be found at http://www.actlabs.com/home.htm. LOI—loss on ignition, TOT—total.
event (Fig. 11D). By this time, the continental deposits of the Rayoso Formation were filling the Neuquén Basin, and basementfault inversion had started in the Agrio fold-and-thrust belt, uplifting the Cerro Mocho anticline in the inner (western) sector (Fig. 7). Shortening between the inner (western) and outer (eastern) sector was accommodated through a basal detachment in the Auquilco evaporites (Figs. 6 and 7) that transferred the horizontal displacement to the east. The outer sector was deformed at this time by thin-skinned tectonics, generating detachment folds seen as the Cerro La Mula, Pichi Mula, etc. Kay (2001) and Kay et al. (this volume, chapter 2) present an 40Ar/ 39Ar biotite cooling age of 69.09 ± 0.13 Ma from a granodiorite pluton near Varvarcó in the western Cordillera del Viento, which they argue marks a time of uplift. Linares and Gonzalez (1990) and Franchini and Schalamuk (1999) presented ages from 65 to 75 Ma for the Paleogene andesitic series
Figure 14. Map of the Collipilli igneous rocks and Cerro Mocho dikes in the framework of the Upper Cretaceous–Paleogene intrusive and extrusive rocks of the northwestern Neuquén province. See text for sources of age dates.
that unconformably overlie the deformed western flank of the Cordillera del Viento. Similar values have been presented here for the igneous rocks of the Collipilli area, which overlie deformed sediments. This event would have produced the uplifting of the Cordillera del Viento as a reactivation of the pre-existing normal faults (Zapata et al., 1999a, 1999b). Late Cretaceous to Early Miocene The structural relief of the Agrio fold-and-thrust belt generated during the first deformational stage could be responsible for the intra–Rayoso Formation unconformity (Ponce et al. 2002), as well as the unconformity between the Rayoso Formation and Neuquén Group. In this sector of the Neuquén Andes,
Structural evolution and magmatic characteristics of the Agrio fold-and-thrust belt the continental sediments of the Neuquén Group were deposited over synclines, such as the Collipilli, Pampa de Naunauco, and those in the external part of the Agrio fold-and-thrust belt that were produced in the first deformation stage. Miocene deformation would have eroded these deposits, leaving only those now found in the outer sector of the Agrio fold-and-thrust belt. At this time, an important magmatic activity started at this latitude with the eruption of the Eocene Cayanta Formation on the back limb of the Cordillera del Viento and intrusion of the laccoliths of Cerro Nevazón, Varvarcó, etc. In the study area, the Collipilli and Naunauco andesitic rocks cut through the Collipilli syncline. The Collipilli and Naunauco magmas have the chemistry expected of magmas erupted in or near the frontal volcanic arc over the subducting Nazca plate. If they are compared with the andesitic series and rocks from Cerro Nevazón (Franchini et al., 2003), their composition, volume, and regional distribution are in accord with eruption through a thin crust through which mantle-generated melts easily passed to the surface. Middle Miocene to Pliocene (Second Compressional Stage) A new change in regional stress produced a new deformational event. In the Agrio fold-and-thrust belt, this event is reflected in the reactivation of basement faults that inverted previous Jurassic half-grabens. It produced the tightening of all previous Cretaceous structures (Fig. 8), such as the Cordon del Salado, Pichi Mula, and Cerro Rayoso anticlines. This uplift is associated with the deposition of synorogenic deposits like those in the Puesto Burgos Formation and the Tralalhué conglomerates (Figs. 2 and 9). This deformation event is also recognized in the Chos Malal area by Kozlowski et al. (1996) where deposits of latest Miocene age and Pliocene are not folded. CONCLUDING REMARKS The new data presented here allow us to create a tectonic model for the history of the Agrio fold-and-thrust belt. The inversion of Jurassic half-grabens structured the inner sector in the latest Early Cretaceous through thick-skinned tectonics. Shortening was transferred to the foreland, producing thinskinned structures in the outer sector. Later, the Collipilli igneous rocks intruded these structures, and a second deformational event took place in the Paleocene to middle Miocene. This event reactivated the previous structures and produced basement-fault inversion in the outer sector. The Tralalhué and Puesto Burgos synorogenic deposits record this event. The Naunauco andesites and Collipilli igneous rocks have geochemical characteristics like those of other Andean-margin arc magmatic rocks erupted through a normal-thickness crust. They are assigned to a redefined latest Cretaceous, MaastrichtianPaleocene Collipilli Group based on Ar/Ar ages ranging from 73 Ma to 65 Ma. The 102 Ma Cerro Mocho volcanic rocks, which were emplaced as E-W–trending dikes, reflect an earlier
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Cretaceous magmatic event that was previously unknown in the region. The Collipilli and Cerro Mocho Groups both have traceelement patterns associated with convergent continental margins, but differences between them require that they came from distinct source regions. REFERENCES CITED Bettini, F.H., 1982, Complejos efusivos Terciarios presentes en las hojas 30c y 32b (Puntilla de Huincán y Chos Malal, del sur de Mendoza y norte de Neuquén), Argentina, in V Congreso Latinoamericano Geológico, Argentina: Actas, V: 79–114 Bracaccini, I.O., 1970, Rasgos tectónicos de las acumulaciones Mesozoicas en las provincias de Mendoza y Neuquén, República Argentina: Asociación Geológica Argentina Revista, v. 25, no. 2, p. 275–282. Brissón, I., and Veiga, R., 1998, La estratigrafía y estructura de la Cuenca Neuquina. Gira de campo: Buenos Aires, Repsol YPF, (unpublished report). Cobbold, P.R., and Rossello, E.A., 2003, Aptian to Recent compressional deformation, foothills of the Neuquén Basin, Argentina: Marine and Petroleum Geology, v. 20, no. 5, p. 429–443, doi: 10.1016/S0264-8172(03)00077-1. Domínguez, E., Aliotta, G., Garrido, M., Daniela, J.C., Ronconi, N., Casé, A.M., and Palacios, M., 1984, Los Maitenes–El Salvaje: Un sistema hidrotermal tipo porfírico, in IX Congreso Geológico Argentino (Bariloche): Actas, v. VII, p. 443–458. Eisner, P., 1991, Tectonostratigraphic evolution of the Neuquén Basin, Argentina [Master’s thesis]: Houston, Rice University, 56 p. Franchini, M., 1992. Las rocas intrusivas del Cerro Caicayén, provincia del Neuquén y su relación con las manifestaciones de hierro en skarns: Revista de la Asociación Geológica Argentina, v. 47, no. 4, p. 399–408. Franchini, M., and Schalamuk, A., 1999, Cuerpos ígneos asociados a skarns en el cinturón plegado de Chos Malal, NO Neuquén: Geoquímica y metalogenesis, in XIV Congreso Geológico Argentino (Salta, Argentina): Actas, v. II, p. 218–221. Franchini, M., López-Escobar, L., Schalamuk, I.B.A., and Meinert, L., 2003, Magmatic characteristics of the Paleocene Cerro Nevazón region and other Cretaceous to Early Tertiary calc-alkaline subvolcanic to plutonic units in the Neuquén Andes, Argentina: Journal of South American Earth Sciences, v. 16, p. 399–421, doi: 10.1016/S0895-9811(03)00103-2. 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. I, 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: Asociación Geológica Argentina, Revista, v. 2, no. 4, p. 347–408. Gulisano, C.A., Gutierrez Pleimling, A.R., 1985, Análisis estratigráfico y sedimentológico de la Formación tordillo en el oeste de la provincia de Neuquén, Cuenca Neuquina, Argentina: YPF, Buenos Aires, unpublished report. Gulisano, C.A., and Gutierrez Pleimling, A.R., 1994, Field guide to the Jurassic of the Neuquén Basin, province of Neuquén: Dirección Nacional del Servicio Geológico, Publicación, v. 158, p. 1–111. Haller, M., Nullo, F.E., Proserpio, C.A., Parica, P.D., and Cagnoni, M.C., 1985, Major element geochemistry of early Tertiary volcanics: Departamento de Geología, Universidad de Chile, Santiago, 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.
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la latitud de 37°30': Dptos de Ñorquin y Pehuenches, Provincia del Neuqén, YPF, unpublished report. Pearce, J., 1996, A user’s guide to basalt discrimination diagrams Geological Association of Canada, Short Course Notes 12, p. 79–114. Ploszkiewicz, J., 1987, Las zonas triangulares de la faja fallada y plegada de la cuenca Neuquina, Argentina, in X Congreso Geológico Argentino: Actas, v. 1, p. 177–180. Ponce J. J., Zavala, C., Marteau, V and Drittanti, D., 2002, Análisis estratigráfico y modelo deposicional para la formación rayoso (cretácico inferior) en la cuenca neuquina, provincia del Neuquén, in Cabaleri, N., Cingolani, C.A., Linares, E., López de Luchi, M.G., Ostera, H.A., and Panarello, H.O., eds., Actas del XV Congreso Geológico Argentino (El Calafate), CD-ROM, Artículo no. 235, 6 p. 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: Boletín del Servicio Geológico Nacional, v. 182, p. 1–103. Ramos, V.A., 1998, Estructura del sector occidental de la faja plegada y corrida del Agrio, cuenca Neuquina, Argentina, in X Congreso Latinoamericano de Geología (Buenos Aires): Actas, v. II, p. 105–110. Ramos, V.A., and Folguera, A., 2006, Tectonic evolution of the Andes of Neuquén: Constraints derived from the magmatic arc and foreland deformation, in Spalletti, L., Veiga, G., Schwarz, E., and Howell, J., eds., The Neuquén Basin: A case study in sequence stratigraphy and basin dynamics: Geological Society of London Special Publication 252, p. 15–35. Ramos, V.A., and Kay, S.M., 1991, Triassic rifting in the Cuyo Basin, central Argentina, in Harmon, R.S., and Rapela, C.W., eds., Andean magmatism and its tectonic setting: Geological Society of America Special Paper 265, p. 79–91. Ramos, V.A., and Kay, S.M., 2006, this volume, Overview of the tectonic evolution of the southern Central Andes of Mendoza and Neuquén (35°–39°S latitude), 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 latitude): Geological Society of America Special Paper 407, doi: 10.1130/2006.2407(01). Rapela, C.W., and Llambías, E.J., 1985, La secuencia andesítica Terciaria de Andacollo, Neuquén, Argentina, in IV Congreso Geológico Chileno (Antofagasta): Actas, v. III, 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., Actas del XV Congreso Geológico Argentino (El Calafate), CD-ROM, Artículo no. 227, 6 p. Rovere, E.I., and Rosello, E., 2001,Evolución geológica durante el Miocene en la región del C° Columpios, 37°S, Andes Neuquinos, Argentina, in XI Congreso Latinoamericano (Montevideo, Uruguay). Sillitoe, R.H., 1977, Permo-Carboniferous, Upper Cretaceous and Miocene porphyry copper-type mineralization in the Argentinian Andes: Economic Geology and the Bulletin of the Society of Economic Geologists, v. 72, p. 99–103. Uliana, M., and Legarreta, L., 1993, Hydrocarbon habitat in a Triassic to Cretaceous sub-Andean setting: Neuquén Basin, Argentina: Journal of Petroleum Geology, v. 16, p. 397–420. 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., et al., eds., Petroleum basins of South America: American Association of Petroleum Geologists (AAPG) Memoir 62, p. 383–402. Viñes, R.F., 1985, Estilos estructurales de la faja plegada occidental Neuquina. Informe preliminary: Buenos Aires, Repsol YPF (unpublished report), 6p., 6 figures. Wichmann, R., 1934, Contribución al conocimiento geológico de los territorios del Neuquén y del Río Negro: Buenos Aires, Ministerio de Agricultura, Dirección General de Minas y Geología, Boletín, v. 39, p. 1–27.
Structural evolution and magmatic characteristics of the Agrio fold-and-thrust belt Yrigoyen, M.F., 1972, Cordillera Principal, in Leanza, A., ed., Geología regional Argentina: Córdoba, Academia Nacional de Ciencias, p. 345–364. Zamora Valcarce, G., Zapata, T., and Del Pino, D., 2005, Edad de la deformación y magmatismo en la faja plegada del Agrio. XVI Congreso Geológico Argentino, La Plata: Actas, v. I, p. 75–78. Zapata, T.R., and Folguera, A., 2006, Tectonic evolution of the Andean fold and thrust belt of the southern Neuquén Basin, Argentina, in Spalletti, L., Veiga, G., Schwarz, E., and Howell, J., eds., The Neuquén Basin: A case study in sequence stratigraphy and basin dynamics: Geological Society of London Special Publication 252, p. 37–56. Zapata, T.R., Brissón, I., and Dzelalija, F., 1999a, The role of basement in the Andean fold and thrust belt of the Neuquén Basin, in McClay, K., ed., Thrust Tectonics 1999: V.I. Londres, Abstracts, p. 122–124. Zapata, T. R., Brissón, I., and Dzelalija, F., 1999b, La estructura de la faja plegada y corrida Andina en relación con el control del basamento de la Cuenca Neuquina: Boletín de Informaciones Petroleras, Tercera Epoca XVI, v. 60, p. 112–121. 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
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estratos Terciarios de la Cuenca Neuquina, Argentina, in V Congreso de exploración y desarrollo de Hidrocarburos (Mar del Plata). Zapata, T.R., Zamora, G., and Ansa, A., 2003, The Agrio fold and thrust belt: Structural analysis and its relationships with the petroleum system Vaca Muerta–Agrio–Troncoso Inferior(!), Argentina: VIII Simposio Bolivariano Tomo, v. I, p. 168–176. Zavala, C.A., 2000, Nuevos avances en la sedimentología y estratigrafía secuencial de la formación Mulichinco en la Cuenca Neuquina: Boletín de Informaciones Petroleras, Tercera Época, v. 65, p. 52–64. Zavala, C.A., Hernán, M., and Mariano, A., 2002, Las facies de la formación Lotena (Jurásico medio) en las áreas de Loncopué y Loma La Lata, Cuenca Neuquina, Argentina, in V Congreso de Exploración y Desarrollo de Hidrocarburos (Mar del Plata, Argentina): Actas. Zollner, W., and Amos, A., 1973, Descripción geológica de la hoja 32b, Chos Malal, Provincia del Neuquén: Boletín del Servicio Nacional de Mineria y Geología, v. 143, p. 1–91.
MANUSCRIPT ACCEPTED BY THE SOCIETY 22 DECEMBER 2005
Printed in the USA
Geological Society of America Special Paper 407 2006
Synrift geometry of the Neuquén Basin in northeastern Neuquén Province, Argentina Ernesto Cristallini* Laboratorio de Modelado Geológico (LaMoGe), Departamento de Ciencias Geológicas, FCEyN, Universidad de Buenos Aires (Consejo Nacional de Investigaciones Científicas y Técnicas), Ciudad Universitaria, Pabellón 2, 1428 Buenos Aires, Argentina Germán Bottesi Alejandro Gavarrino Leonardo Rodríguez Repsol YPF, Diagonal Roque Saenz Peña 777, (1035), Buenos Aires, Argentina Renata Tomezzoli Instituto de Geomagnetismo “Daniel Valencio,” Departamento de Ciencias Geológicas, FCEyN, Universidad de Buenos Aires (Consejo Nacional de Investigaciones Científicas y Técnicas), Ciudad Universitaria, Pabellón 2, 1428 Buenos Aires, Argentina Raúl Comeron Repsol YPF, Diagonal Roque Saenz Peña 777, (1035), Buenos Aires, Argentina
ABSTRACT A map of the principal grabens and half-grabens of the Neuquén Basin in the northeastern part of Neuquén Province, Argentina, was constructed based on twodimensional (2-D) and three-dimensional (3-D) seismic information. The synrift structure of the region is characterized by high angle NW-trending normal faults. Two populations can be distinguished: one with an azimuth of 140° and the other one with an azimuth of 105°. The principal half-grabens are accommodated by NEdipping normal faults. The principal NW-trending faults are less than 20 km long and are either crosscut by or terminate in transfer zones trending to the NE. The transfer zones are either faults or zones where the principal NW-trending faults loose slip or terminate. The principal normal faults were active until the deposition of the Lower Upper Jurassic Tordillo Formation. Subsequently, only a few faults related to differential subsidence over the half-grabens remained active. The typical structures of the region are smooth anticlines and synclines that affect the sag facies of Neuquén Basin. The anticlines developed over basement highs and the synclines over graben basins. The synclines are explained in this study as resulting from differential subsidence over the half-grabens. This differential subsidence could have been a continuous process that began in the synrift stage of the basin. A distinct element model was
*E-mail:
[email protected].
Cristallini, E., Bottesi, G., Gavarrino, A., Rodríguez, L., Tomezzoli, R., and Comeron, R., 2006, Synrift geometry of the Neuquén Basin in northeastern Neuquén Province, Argentina, 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. 147–161, doi: 10.1130/2006.2407(07). For permission to copy, contact
[email protected]. ©2006 Geological Society of America. All rights reserved.
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E. Cristallini et al. created to analyze the differential subsidence process and geometry. No evidence was found to support the hypothesis that tectonic inversion is an important process in this region. The main unconformities can be explained by the differential subsidence process. Keywords: Neuquén Basin, differential subsidence, rifting, Entre Lomas Height.
INTRODUCTION The Neuquén Basin is a retroarc rift-related basin in southwestern South America that has developed since Triassic times (Fig. 1). Synrift sequences were deposited until the middle of the Jurassic and were followed by typical sag phase sequences, which were deposited up until Upper Cretaceous time. Tectonic inversion associated with the Andean orogeny began in the Upper Cretaceous on the western side of the basin, with continental sediments accumulating in the foreland. Earlier episodes of inversion in the Jurassic in the southern part of the basin can be related to strike-slip and wrench tectonics (Dorsal de Huincul area; Orchuela et al., 1981; Ploszkiewicz et al., 1984; Bettini, 1984). The Neuquén Basin is the most important Argentine basin in terms of hydrocarbon production, accounting for nearly
half of the oil and gas produced in the country. In this paper, we analyze the northeastern region of the basin, which is one of the most important areas for oil and gas production. The northeastern border of the basin developed over a first-order crustal discontinuity (Dalmayrac et al., 1980; Ramos 1984, 1988; Tomezzoli, 2001). This discontinuity controlled the Triassic synrift structures in the northeastern part of the basin (Ramos, 1984). The oil fields of this area have an important structural control (Veiga et al., 1999) related to the principal faults that bound the Triassic half-grabens. These NW-trending faults define anticlinal highs and synclinal troughs in all of the post-Triassic sedimentary sequences. These structural highs and lows also exert a control on the modern topography and drainage. An important question is why the Triassic structures control the subsequent
Figure 1. Location map of the Neuquén Basin and the neighboring geological provinces of southern Argentina. The area considered in this study is shown in the boxed region, which corresponds to the region shown in Figure 2.
Synrift geometry of the Neuquén Basin sedimentary history of the region. Is tectonic inversion controlling the structure in this region, as is the case in the western and southern parts of the Neuquén Basin? Is extensional reactivation of Triassic structures playing a role? Or is the main control differential subsidence? In this paper, we present an integrated map based on two-dimensional (2-D) and three-dimensional (3-D) seismic information that shows the arrangement and geometry of the principal Triassic half-graben structures of the northeastern Neuquén Basin. We then analyze the mechanism by which the basement structures are controlling the sedimentation and structure of the uppermost portion of the sedimentary cover.
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OBSERVATIONS Map of Principal Synrift Faults A main objective in this investigation was the construction of a regional map of the principal grabens and half-grabens of the Neuquén Basin in the northeastern part of Neuquén Province. The principal data used were several 2-D seismic lines from the Repsol YPF company. Some 3-D seismic and borehole information were used as a check in specific regions. The map of the principal graben and half-graben structures is shown in Figure 2. The faults are categorized according to importance, with their relative importance indicated by the
Figure 2. Map of the principal grabens and half-grabens of the region. Colors represent seismic thickness measured in two-way travel time (TWT) units (ms) of pre-Cuyo Group deposits. RCH/VE—Rincon Chico–Veta Escondida lineament. The northern (NH), central (CH) and southern (SH) highs discussed in the text are shown. Locations of the cross-sections A and B are indicated. The trace thickness used to represent faults and hinges is proportional to the half-graben and fault importance (see reference on text). Map is based on the Gauss-Kruger coordinate system (Argentina-zone 2) using the Campo Inschauspe Datum; units are meters.
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thickness of the fault trace. The importance of a fault is a subjective characteristic and can change during the history of the fault. In this paper, we consider the importance of each fault only for the synrift stage of the basin. The variables used to calculate the importance of a fault are: (1) In cases where the fault represents the boundary of a graben or half-graben, the thickness of synrift deposits was considered as the most sensitive parameter in determining importance. (2) In cases where the fault does not represent the boundary of a graben or half-graben (i.e., internal faults of a half-graben), the slip of the fault was considered as the most sensitive parameter in determining importance. (3) A subjective value was assigned based on the aspect of the fault in the seismic line. The following equation was used in the calculation of the importance of each fault: I = (T × KT + S × KS)V,
(1)
where I is the importance of the fault, T is the parameter representing the thickness of the graben or half-graben (with values from 1 to 8), S is the parameter representing the slip of the fault (with values from 1 to 8), V is the subjective parameter assigned according to the aspect of the fault in the seismic lines (with values from 1 to 3), and KT and KS are Boolean variables. The value of KT is 1 in number (1) situations and 0 in number (2) situations, and the value of KS is 1 in number (2) situations and 0 in number (1) situations. The thickness of the fault traces in Figure 2 is proportional to I. The summary map in Figure 2 shows that the synrift faults have high dip angles and generally trend to the NW. Two populations are present, one with an average azimuth of 140° and the other with an average azimuth of 105°. Overall, 52% of the analyzed faults dip to the SW and 48% dip to the NE. The most important half-graben system in the region is located just to the northeast of the Entre Lomas high and is bounded by NEdipping faults. There is no relation between trend populations and dip directions. Transfer zones with NE trends were also identified and mapped in Figure 2. The transfer zones are defined by aligned sectors where the main faults lose continuity. In a few cases, they are represented by NE-trending faults. One of these transfer zones coincides with the Rincón Chico–Veta Escondida alignment described by Veiga et al. (1999). Transfer zones can exert an important control on surface features, as shown by the position of the Auca Mahuída volcanic complex and some of the sharp curves of the Colorado River. The map in Figure 2 shows half-graben configurations at the end of the pre-Cuyo Group synrift in the Lower Jurassic. Thicknesses of synrift deposits are denoted by the corresponding seismic (two-way travel) time interval (isocronopachic map). Three main highs are observed and designated as the northern high, the southern high, and the central high. The northern high (NH in Fig. 2) is an 8-km-wide and 100-km-long horst located near the northeast boundary of the Neuquén Basin. It is a symmetric NW-
trending structure bounded by NE-dipping faults to the north and SW-dipping faults to the south. The southern high (SH in Fig. 2) is equivalent to the Entre Lomas–Señal Cerro Bayo (Veiga et al., 1999) lineament and corresponds to a NW-trending horst. It is a clear asymmetric structure where the main fault bounds the structure to the north and dips northeast. The horst is limited to the south by a less important, SW-dipping fault. Compared to the northern high, the southern high is similar in length, narrower, and more fragmented by transfer zones. Compared to the other two highs, the central high (CH in Fig. 2) is less important and more complex, in that it is more segmented and branched out. In general, this high is formed by the elevated parts of half-graben hanging walls, and some segments have a horst structure. The principal synrift basin of the region is located between the central and southern highs, and the deepest and bestdefined troughs occur in this area. The troughs are asymmetric basins (half-grabens) that have principal faults trending SW and dipping NE. Structural Sections Two cross sections were constructed using 2-D seismic information to analyze and illustrate the structure of the region. Their locations are shown in Figure 2. Cross-section A in Figure 3 cuts the northern, central, and southern structural highs discussed above. The half-grabens between the highs are asymmetric, with their thickest parts generally adjacent to the principal faults to the southwest. The most important half-graben occurs between the southern and central highs. The postrift units (e.g., Quintuco and Centenario Formations) show a regional thickening to the southwest toward the center of the basin, and in some cases show a distortion related to the synrift faults. In this section, it is clear that the uppermost units of the sequence have a synclinal curvature just above the corresponding half-graben. This is common throughout the region and can be utilized to trace the axis of the principal halfgrabens, even when deep seismic information is poorly defined. Cross-section B in Figure 3 crosses the southern and central highs further to the northwest (Fig. 2). As in section A, the principal half-graben occurs between the southern and central highs, the sag and foreland units show a regional thickening to the southwest toward the center of the basin, and a local collapse occurs above each half-graben. The central high seems to act as a hinge between the principal grabens to the northeast and southwest. Timing of Deformation All the faults on the map in Figure 2 show evidence for important activity during the synrift stage (pre-Cuyo Group). They controlled the shape and position of the principal Triassic half-grabens. Because later movements on these faults are very small and difficult to analyze using 2-D seismic information, 3-D seismic lines were employed to study post-Triassic activity on the principal faults.
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Figure 3. Cross sections of the region constructed using two-dimensional (2-D) seismic information. See location in Figure 2. TWT—two-way traveltime.
As an example, Figure 4 shows a 3-D seismic line from Bajo del Piche (see Fig. 2 for location), where three principal faults, labeled BP1, BP2, and BP3, can be seen. The first, BP1, is one of the major faults of the region. It bounds the northeastern side of the southern high (Entre Lomas high) and lies southwest of the principal half-graben in the region. This fault was only active during the synrift stage. In contrast, fault BP2, on the southwestern limit of the southern high, was not as important during the synrift period, but continued to move until at least the Cretaceous. The principal motion on this fault occurred during the Triassic and then diminished, becoming smaller during the Upper Jurassic (up to the time of the Tordillo Formation) and even smaller during the Cretaceous (up to time of the Rayoso Formation). The timing information from all of the faults that were analyzed in the region is illustrated in Figure 5. The four maps sequentially show faults that were active up until the Late Jurassic (Tordillo Formation), the Lower to Middle Cretaceous (Centenario Formation), the Upper Cretaceous (Neuquén Group), and those that are affecting the present surface. Comparing Figures 2 and 5A, one can see that almost all the synrift faults developed during the Triassic were active up until the end of the deposition of the Tordillo Formation in the Late Jurassic. This pattern is also
supported by well information. As an example, Figure 6 shows the thickness of the principal units in two contiguous wells, one over a horst and the second over a half-graben (see Fig. 2 for location). The thicknesses of the Lower to Middle Cretaceous units, above the Late Jurassic Tordillo Formation, increase to the south following the regional slope of the basin. In contrast, the thickness of the Tordillo Formation increases to the north, following the local slope of the half-graben and showing that this half-graben was active in the Late Jurassic. INTERPRETATION Deformation after Synrift Stage The origin of the anticlines that affect the postrift units in this part of the Neuquén Basin is considered below. An important observation is that the postrift units above the highs or anticlines are generally thicker than those in the troughs or synclines to the northeast, indicating that the deformation must be postsedimentation. Three possible deformational mechanisms that may explain the present configuration are shown in Figure 7: tectonic inversion, normal fault reactivation, and differential subsidence related to differential compaction. The main problem
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Figure 4. Structure of Bajo del Piche based on a three-dimensional (3-D) seismic line. Principal faults BP1, BP2, and BP3 are indicated. See Figure 2 for location.
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Figure 5. Maps showing the faults that report activity until the Late Jurassic (Tordillo Formation), Lower to Middle Cretaceous (Centenario Formation), Upper Cretaceous (Neuquén Group), and faults that are affecting the present surface. Red traces indicate fault activity, green traces indicate fold activity, and blue traces indicate no activity in the period. Each map represents the same area and structure of Figure 2.
is to establish the regional level of a unit (level of the unit before deformation). In the first case in Figure 7A, the regional level is between the hinges of the synclines. In the cases in Figures 7B and 7C, the regional level must be in the hinge of the anticline. Figure 8 shows theoretical separation diagrams for the three mechanisms illustrated in Figure 7. These diagrams show plots of the separation from the regional level that a postrift unit acquires during deformation (dashed line) compared to the original shape of the synrift unit (gray line). For the case analyzed here, we constructed a separation diagram based on the principal units in crosssection A in Figure 3. A tracing of the principal units is shown in Figure 9. The regional level for each unit is arbitrarily placed in the anticline hinges (PC—pre-Cuyo Group, T—Tordillo Formation, Q—Quintuco Formation, C—Centenario Formation).
The central part of the section was selected to make the separation analysis. Figure 10 shows the resulting diagram for all of the markers; the thickness of the entire synrift unit is shown as a reference. A direct relation between the separation from regional levels of the postrift units and synrift thickness can be observed. A comparison between Figure 10 and the theoretical diagrams in Figure 8 shows that the separation curves in Figure 10 are most like the differential subsidence case in Figure 8A. The model for tectonic inversion can be dismissed, since Figure 10 shows exactly the opposite of what is seen in Figure 8C, and no evidence for tectonic inversion was observed in the seismic lines. The reactivation of normal faults model cannot be rejected based on these analyses, since the separation diagram in Figure 8B is similar to that for differen-
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Figure 6. Variation of the thickness of the principal units in two contiguous wells: 1 and 2. Well 1 is over a horst and is south of well 2, which is over a half-graben. Note that the thickness of the Tordillo Formation increases to the north, following the local slope of a half-graben, whereas the thickness of the overlying Lower to Middle Cretaceous units increases to the south, following the regional slope of the basin. See Figure 2 for location of wells.
tial subsidence in Figure 8A, and there is evidence for fault reactivation all over the region. As seen in Figures 8A and 8B, the main difference in the theoretical diagrams for differential subsidence and reactivation of normal faults is that the internal shape of the halfgraben is followed by the separation curve in the differential subsidence case. In the case of normal fault reactivation, only the boundaries of the half-graben control the shape of the separation curve. In Figure 10, we can see that the shape of the separation curves mimics the shape of the half-grabens, especially the main one. As such, differential subsidence is the most probable mechanism to explain structural highs seen in the upper units. Figure 6 shows that the thicknesses of units above the Tordillo Formation are greater in the high (horst) and smaller in the trough (graben) to the northeast. This is a common case for the uppermost units in other parts of the region, implying that differential subsidence took place after the formation of these units. This observation creates a problem for the differential subsidence model, because it is more intuitive to think that dif-
ferential subsidence acts continuously. In the following, we show how this thickness relation between the highs and troughs can be explained using a continuous mechanism. Theoretical Simulation of Differential Subsidence A qualitative numerical simulation of the differential subsidence mechanism was preformed for cross-section A using the discrete element model of Cundall and Starck (1979) implemented by Zlotnik (2002) software AsBolinhas®. In the model, the rocks are simulated by individual circular elements that are separated from each other by frictional surfaces. A spring concept is used to join the individual elements to simulate a cohesive material. Each element and each spring is given physical properties (size, density, stiffness, etc.), and a statistical combination of these characteristics gives the material physical properties (density, Young modulus, Poisson coefficient, cohesion, internal friction, etc.). Two simplified simulations of the half-grabens were performed, one without springs (noncohesive material; Fig. 11)
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Figure 7. Three different mechanisms of deformation to arrive at the present configuration were evaluated: tectonic inversion, normal fault reactivation, and differential subsidence. RL—regional level.
Figure 8. Theoretical separation diagrams are shown for the three mechanisms evaluated. These diagrams plot the separation from regional level that a postrift unit acquires during deformation (dashed line) in comparison to the shape of a synrift unit (gray line).
and the other with springs (cohesive material; Fig. 12). In both, schematic half-grabens were separated by a structural high that represents the Entre Lomas high in section A (Fig. 3). The basement was not allowed to deform, and the faults that bound the half-grabens were constrained not to move during the simulation. Both models had a similar design with 36,000 elements, and each ran for several days on a personal computer. Differential subsidence occurred in both models.
Figure 11A shows the results of the noncohesive simulation, in which anticlines of postrift material developed over the structural highs. Figure 11B shows the separation diagram for one layer in the model. A distinctive feature of the model is that the separation from the regional level occurred smoothly over the structural highs. Figure 12A shows the results of the cohesive simulation. In contrast to the noncohesive simulation, faults grew upward from the original synrift boundaries and cut
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Figure 9. Tracing of the principal stratigraphic units in cross-section A in Figure 3. The regional level for each unit is arbitrarily placed in the anticlinal hinges. Abbreviations: PC—pre-Cuyo Group, T—Tordillo Formation, Q—Quintuco Formation, C—Centenario Formation.
Figure 10. Resulting separation diagram for all the markers shown in Figure 9; the thickness of synrift unit is indicated as a reference. The scale on the left shows thickness, measured by two-way traveltime (TWT) in seconds, on a seismic line. The separation scale on the right shows the thickness transformed to distance using a regional velocity law.
the lowermost sag sequence. The diagram in Figure 12B shows an abrupt separation just above the synrift faults. The elements that broke springs during the simulation are shown in black in Figure 12A to emphasize the upward fault propagation that occurred in the synrift faults. The separation diagram in Figure 10, based on cross-section A (Fig. 3), is more similar to that for the cohesive simulation in Figure 12 than for the noncohesive simulation in Figure 11. The analysis shows that some of the faults that cut the sag faces can be explained by fault propagation due to differential subsidence alone. No extensional tectonic mechanism is needed.
Evaluation of Differential Subsidence Mechanism In order to evaluate the viability of differential subsidence as a generalized mechanism for the region, we decompacted the fill of the half-graben to the northeast of the Entre Lomas high. The shape of the half-graben is compared to three different decompaction curves for the top of the Tordillo Formation in Figure 13. The three curves are based on different porosity and compaction coefficients. Assuming a fill of 70% sandstone and 30% shale, the shape of the top of the Tordillo Formation marker is most easily restored to a planar outline using a porosity
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Figure 11. (A) Differential subsidence simulation with noncohesive material using distinct elements method. Anticlines of postrift units develop above the structural highs. (B) The separation diagram of a selected layer of the model.
Figure 12. (A) Differential subsidence model using distinct elements method. Cohesive material was simulated using springs between elements. Anticlines of postrift units develop above the structural highs. (B) Separation diagram of a selected layer of the model. (C) Detail of the model where the elements that broke springs are blacked to illustrate fracturing and faulting nucleated above basement faults.
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Figure 13. The shape of the half-graben to the northeast of the Entre Lomas high and curves of decompaction for the top of the Tordillo Formation for different porosity and compaction (c) coefficients (see reference in text).
of 53% and a compaction coefficient of 0.34 km–1. However, the main purpose here is not to match the porosity and compaction coefficients of the half-graben filling, but to show that common synrift lithologies can produce differential compaction that is consistent with differential subsidence. Since the infilling of the pre-Cuyo half-grabens contains volcaniclastic rocks in some regions, we have done calculations that show that the change in volume produced by alternation of volcanic glass also contributes to differential compaction. It is reasonable to think that differential compaction and subsidence should be linked in a continuous mechanism. A problem with such a linkage in the region of interest is that the uppermost postrift units (generally post-Tordillo Formation) are generally thicker over the highs than they are in adjacent troughs to the northeast. In other words, the postrift thicknesses are mainly controlled by regional slope, not by differential subsidence. As a first approximation, this relation indicates that differential subsidence occurred after deposition and did not control sedimentation rates. A further analysis shows that the difference between the relative velocity of regional subsidence due to thermal effects and the relative velocity of local subsidence due to differential compaction controls the relative thicknesses of the sediments over the highs and in the troughs. Three cases are shown in Figure 14: (A) only regional subsidence; (B) differential subsidence slower than regional subsidence; and (C) differential
subsidence faster than regional subsidence. The relative sediment thicknesses shown by the markers in Figure 14 demonstrate that when regional subsidence is faster than local subsidence, thickness increases toward the interior of the basin, and that when the reverse occurs, thickness increases in the troughs and decreases in the highs. In the cases where differential subsidence is occurring (Figs. 14B and 14C), smooth anticlines develop over the basement highs and smooth synclines over the basement troughs. The case in Figure 14B can be used to explain the geometry and sedimentation of the postrift units above the Tordillo Formation, and the case in Figure 14C, to explain those of the lower sag units up to the Tordillo Formation. In this way, we can explain the geometries and sedimentation pattern of the postrift units in the study region by continuous differential subsidence. The difference between the discrete reactivation of normal faults and continuous differential subsidence models is significant for the evolution of the region. Given a differential subsidence model, basement structure exerts a control on sedimentation on all units above the basement, as well as the distribution of Mesozoic to Holocene fluvial paths. Basement control in the sedimentation of postrift units can be clearly seen until the Upper Jurassic (Tordillo Formation). We can see this by comparing the paleogeographic reconstruction of Cazau and Melli (2002) for the Sierras Blancas Formation (Tordillo Formation equivalent) with the synrift structural map in Figure 2 and noting that the
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Figure 14. Relative rates of regional and local subsidence. Three cases are shown: (A) there is no differential subsidence (local subsidence); only regional subsidence is acting; (B) the differential subsidence is slower than the regional subsidence; and (C) the differential subsidence is faster than the regional subsidence. Numbers indicate thicknesses between highlighted layers in different positions.
principal rivers either parallel principal faults or transfer zones (Fig. 15). 3-D seismic data provide clear evidence for such control up until the Lower Cretaceous (Quintuco Formation). Later, the situation is less clear, because seismic evidence is poor. Nevertheless, evidence for basement control comes from part of the present trace of the Colorado River and other modern fluvial courses that follow principal alignments of synrift faults or transfer zones. CONCLUSIONS We have analyzed the synrift structure of the northeastern region of the Neuquén Basin and produced an integrated regional map showing the location of the half-graben faults and transfer zones. The faults show a NW trend and generally dip to the NE or
SW. Those that accommodate the main half-grabens are mostly NE-dipping. The transfer zones trend to the NE and divide the principal faults into segments that are shorter than 20 km. The half-grabens and principal faults that affect the region developed in the synrift stage in the Upper Triassic to Lower Jurassic. During the subsequent sag stage, the activity of the principal synrift faults controlled sedimentation until at least the deposition of the Tordillo Formation in the Upper Jurassic. We suggest that this activity was due to differential subsidence related to differences in compaction coefficients between the synrift deposits and basements rocks. In the Lower Jurassic, the velocity of local subsidence due to compaction was higher than the velocity of regional thermal subsidence. This relation subsequently inverted, and regional subsidence became faster. We did not find any evidence to support tectonic inversion
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Figure 15. Paleogeographic reconstruction of Sierras Blancas Formation (Cazau and Melli, 2002) superimposed on the synrift structural map in Figure 2. Note that the principal fluvial courses are parallel to either the principal faults or the transfer zones. Sediment accumulation areas are shown in yellow and emergent areas in gray. Rivers were parallel to yellow trends.
as the mechanism that produced the anticlines and synclines in the northeastern region of the Neuquén Basin. In contrast, tectonic inversion related to transpression in the Jurassic played an important role in developing the structures to the south and southwest, in the region of the Huincul Ridge (Dorsal Neuquina) (Veiga et al., 1999; Mosquera, 2002; Pángaro et al., 2002). Similarly, tectonic inversion that initiated in the Upper Cretaceous played an important role in the development of the fold-andthrust belt in western Neuquén and south Mendoza (Zapata et al., 1999; Kozlowski 1996; Manceda and Figueroa, 1995). The differences between the study area and the other regions can be reconciled by their relative locations (see Fig. 1). The typical structure of the study region in the northeastern region of the Neuquén Basin consists of anticlinal highs and synclinal troughs that developed in postrift sedimentary
sequences. We propose that a relatively simple differential subsidence model can explain their origin. No regional extension is needed to reactivate old normal faults or produce new ones. All of the structures in the Mesozoic to early Tertiary sequences can be explained by continuous subsidence controlled by differential compaction between synrift deposits and basement rocks. ACKNOWLEDGMENTS This work was founded by Repsol YPF, Universidad de Buenos Aires, Fundación Antorchas grants to Ernesto Cristallini and Renata Tomezzoli, and Consejo Nacional de Investigaciones Científicas y Técnicas grants to Ernesto Cristallini and Renata Tomezzoli. Special thanks are due to Estanislao Kozlowski and an anonymous reviewer.
Synrift geometry of the Neuquén Basin REFERENCES CITED Bettini, F.H., 1984, Pautas sobre cronología estructural en el área del Cerro Lotena, Cerro Granito y su implicancia en el significado de la Dorsal del Neuquén, Provincia del Neuquén, in Noveno Congreso Geológico Argentino: Actas, v. 2, p. 342–361. Cazau, L., and Melli, A., 2002, La Formación Sierras Blancas en el noreste de la Cuenca Neuquina, in V Congreso de Exploración y Desarrollo de Hidrocarburos (Mar del Plata): electronic publication, CD-ROM. Cundall, P.A., and Starck, D.L., 1979, A discrete numerical model for granular assemblies: Geotechnique, v. 29, no. 1, p. 47–65. Dalmayrac, B., Laubacher, G., Marocco, R., Martinez, C., and Tomasi, P., 1980, La chaine Hercynienne d’Amérique du Sud: Structure et évolution d’un orogène intracratonique. Sonderdruck a.d.: Geologische Rundschau, v. 69, no. 1, p. 1–21, doi: 10.1007/BF01869020. Kozlowski, E., 1996, Geología estructural de la zona de Chos Malal, Cuenca Neuquina, Argentina, in XIII Congreso Geológico Argentino y III Congreso de Exploración de Hidrocarburos: Actas, v. 1 p. 15–26. Manceda, R., and Figueroa, D., 1995, Inversion of the Mesozoic Neuquén rift in the Malargüe fold-thrust belt, Mendoza, Argentina, in Tankard A.J., Suarez S., R., and Welsink, H.J. eds., Petroleum basins of South America: American Association of Petroleum Geologists Memoir 62, p. 369–382. Mosquera, A., 2002, Inversión tectónica Jurásico inferior en el sector central de la dorsal de Huincul, área Los Bastos, in V Congreso de Exploración y Desarrollo de Hidrocarburos (Mar del Plata): electronic publication, CD-ROM. Orchuela, I.A., Ploszkiewicz, J.V., and Viñes, R.F., 1981, Reinterpretación structural de la denominada “Dorsal de Huincul,” in Octavo Congreso Geológico Argentino: Actas, v. 3, p. 281–293.
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Pángaro, F., Veiga, R., and Vergani, G., 2002, Evolución tecto-sedimentaria del área de Cerro Bandera, Cuenca Neuquina, Argentina, in V Congreso de Exploración y Desarrollo de Hidrocarburos (Mar del Plata): electronic publication, CD-ROM. Ploszkiewicz, J.V., Orchuela, I.A., Vaillard, J.C., and Viñes, R.F., 1984, Compresión y desplazamiento lateral en la zona de Falla de Huincul, estructuras asociadas, Provincia del Neuquén, in Noveno Congreso Geológico Argentino (Bariloche, Buenos Aires): Actas, v. 2, p. 163–169. Ramos, V.A., 1984, Patagonia: Un continente Paleozoico a la deriva?, in Noveno Congreso Geológico Argentino (Bariloche, Buenos Aires): Actas, v. 2, p. 311–325. Ramos, V.A., 1988, Tectonics of the Late Proterozoic–early Paleozoic: A collisional history of southern South America: Episodes, v. 11, no. 3, p. 168–174. Tomezzoli, R.N., 2001, Further palaeomagnetic results from the Sierras Australes fold and thrust belt, Argentina: Geophysical Journal International, v. 147, p. 356–366, doi: 10.1046/j.0956-540x.2001.01536.x. Veiga, R., Lara, M.E., and Bruveris, P., 1999, Distribución de hidrocarburos sobre el margen externo en una cuenca de tras-arco. Ejemplos en la cuenca Neuquina, Argentina: Boletín de Informaciones Petroleras (BIP), v. 60 (Diciembre 1999), p. 142–164. Zapata, T.R., Brissón, I., and Dzelalija, F., 1999, La estructura de la faja plegada y corrida Andina en relación con el control del basamento de la cuenca Neuquina: Boletín de Informaciones Petroleras (BIP), v. 60, p. 112–121. Zlotnik, S., 2002, Modelado numérico de estructuras y deformación de rocas mediante el método de elementos discretos [Trabajo Final de Licenciatura]: Buenos Aires, Universidad de Buenos Aires, 70 p.
MANUSCRIPT ACCEPTED BY THE SOCIETY 22 DECEMBER 2005
Printed in the USA
Geological Society of America Special Paper 407 2006
The case for extensional tectonics in the Oligocene-Miocene Southern Andes as recorded in the Cura Mallín basin (36°–38°S) W. Matthew Burns* Teresa E. Jordan Department of Earth and Atmospheric Sciences, Cornell University, Ithaca, New York 14853-1504, USA Peter Copeland Department of Geosciences, University of Houston, Houston, Texas 77204, USA Shari A. Kelley Department of Earth and Environmental Science, New Mexico Institute of Mining and Technology, Socorro, New Mexico 87801, USA
ABSTRACT The Cura Mallín basin is part of a chain of sedimentary basins that formed within the Andean volcanic arc between 33° and 43°S during the late Oligocene and early Miocene. Most previous studies of these basins have suggested that they are pull-apart–type basins, produced by strike-slip deformation of the Liquiñe-Ofqui fault zone and other structures, all of which are currently active. However, no direct evidence has been cited for a correlation between formation of the Oligocene-Miocene basins and concurrent strike-slip faulting. The Cura Mallín basin lies more than 100 km north of the modern Liquiñe-Ofqui fault zone and is one of the largest and best exposed of the Southern Andean Oligocene-Miocene basins, making it a promising study area for distinguishing between Oligocene-Miocene tectonic activity that produced the basin and subsequent tectonic activity. Stratigraphic and structural data presented here from the Cura Mallín basin and its surroundings include facies variations, stratal thickness patterns, internal and external structural features, 40Ar/ 39Ar radiometric ages, and apatite and zircon fission-track ages. Based on the distribution of sedimentary facies and their relation to geologic structures, we conclude that the Cura Mallín basin formed as a result of normal faulting, with little or no significant strike-slip deformation in the area. Due to the lack of supporting evidence for interpretations of the other Oligocene-Miocene basins as pull-apart basins, we suggest that the entire chain of Oligocene-Miocene sedimentary basins formed in response to extensional tectonics on the Southern Andean margin. Keywords: Andes, Oligocene, Miocene, extension, basin analysis.
*Current address: U.S. Geological Survey, Reston, Virginia 20192, USA; e-mail:
[email protected].
Burns, W.M., Jordan, T.E., Copeland, P., and Kelley, S.A., The case for extensional tectonics in the Oligocene-Miocene Southern Andes as recorded in the 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, doi: 10.1130/2006.2407(08). For permission to copy, contact
[email protected]. ©2006 Geological Society of America. All rights reserved.
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INTRODUCTION In the late Oligocene and early to middle Miocene, a chain of sedimentary basins formed between ~33° and 43°S in what are now the Southern Andes and the Chilean Central Valley (Fig. 1). The basins are best known within the modern Southern Andes where they outcrop extensively (Niemeyer and Muñoz, 1983; Muñoz and Niemeyer, 1984; Suárez and Emparan, 1995; Spalletti and Dalla Salda, 1996; Charrier et al., 1996, 2002; Jordan et al., 2001), but their importance in the subsurface of the Central Valley is confirmed by boreholes and by geophysical methods (McDonough et al., 1997; Vergara et al., 1997). The basins formed within and around the contemporary volcanic arc and contain up to 4 km of mostly volcaniclastic continental strata with lava flows and subvolcanic intrusive rocks, lacustrine limestones, and, in the south, marine strata. Although different formation names have been applied locally, at a regional scale there is considerable similarity of depositional environments and compositions. In some sectors, the volcaniclastic strata span the late Eocene through middle Miocene, and in the Lonquimay region, the basin is exclusively of middle Miocene age (Suárez and Emparan, 1995; Charrier et al., 2002). However, most subbasins are filled primarily with strata of late Oligocene to early Miocene age (e.g., Spalletti and Dalla Salda, 1996; Vergara et al., 1997; Jordan et al., 2001), and our study focuses on this particular span of time. The pervasiveness of basin formation over such a large portion of the Andean continental margin in the late Oligocene to early Miocene and the similarities of the basins suggest that a common mechanism, operating on a regional scale, may have produced all of the basins (Jordan et al., 2001). In studies of many of these basins, it has been proposed that they are pullapart–type basins formed by coeval motions of hypothetical and documented strike-slip faults (Suárez and Emparan, 1995; Spalletti and Dalla Salda, 1996; Godoy et al., 1999; Folguera et al., 2002; Kay and Mpodozis, 2002). From the perspective of the south end of the basin system (i.e., south of 39°S), strike-slip deformation of the Liquiñe-Ofqui fault zone (Cembrano et al., 1996; Diraison et al., 1998) encourages the supposition of a transtensional setting for the Oligocene-Miocene basins. However, no direct evidence has been cited for a causal relationship between the strike-slip faults and any of the basins, and current data on the Liquiñe-Ofqui fault zone (Diraison et al., 1998) indicate that the strike-slip system was not active during the main period of basin formation, i.e., the late Oligocene and early Miocene. Therefore, the spatial coincidence of the OligoceneMiocene basins and strike-slip deformation may simply represent superposition of the effects of subsequent tectonic events on basins formed within a very different tectonic setting. The Cura Mallín basin (36–38°S) appears to be one of the most promising study areas within the chain of OligoceneMiocene basins for distinguishing between Oligocene-Miocene tectonic activity that initiated basin formation and subsequent tectonic activity that distorted the basin. The northernmost por-
Figure 1. Map of Chile and western Argentina between 33°S and 43°S latitude. General areas of Eocene to Miocene basins mentioned in the text are shown, as is the Liquiñe-Ofqui fault zone, and the relative positions of the Main Cordillera, the Central Valley, and the Coastal Cordillera. The heavy, dashed line indicates the international boundary between Chile and Argentina.
The case for extensional tectonics in the Oligocene-Miocene Southern Andes tion of the Liquiñe-Ofqui fault zone lies more than 100 km to the south of the basin (Fig. 1), and only a single phase of deformation, a late Miocene compressional inversion of the basin fill, overprints the original basin (Jordan et al., 2001; this study). Paying specific attention to the implications for basin formation, we present here the results of stratigraphic and structural analyses in and around the eastern Cura Mallín basin. The data examined include facies variations, stratal thickness patterns, internal and external structural features on the basis of field mapping, 40Ar/ 39Ar isotopic ages, and apatite and zircon fissiontrack ages. Our preferred interpretation of these data is that the basin formed as a result of normal faulting within a broad volcanically active region, with no confirmed significant strike-slip deformation in the area. LOCATION OF THE STUDY AREA The Cura Mallín basin lies between ~36° and 38°S latitude (Fig. 1) and is exposed at the surface of much of the main Andean Cordillera in Chile and Argentina (Niemeyer and Muñoz, 1983; Muñoz and Niemeyer, 1984) where it has been uplifted and deformed by E-W compression. However, late Miocene to Holocene volcanism covered much of the basin area subsequent to compressional uplift and limits exposure of the basin fill to the deeper river valleys on the Argentine (eastern) side of the basin (Fig. 2). On the Chilean (western) side, Cura Mallín strata comprise a larger percentage of the surficial geology, but access is inhibited by vegetation. The basin also spans much of the Chilean Central Valley (Elgueta and Rubio, 1991; Vergara et al., 1997), although generally in the subsurface, for a total E-W extent approaching 200 km. Lateral transitions between the Cura Mallín basin and the Coya-Machalí and Lonquimay basins, to the north and south, respectively, are poorly understood because intervening zones are covered by young volcanic rocks. Chronologic data confirm that the Cura Mallín Formation correlates with the upper, most voluminous portion of the Coya-Machalí Formation north of 36°S (Charrier et al., 2002). However, the Oligocene–lower Miocene strata of the eastern Cura Mallín basin, treated here, may be largely older than the strata of the Lonquimay basin south of 38°S (Suárez and Emparan, 1995). Our work focused on the eastern portion of the Cura Mallín basin in Argentina, where it is exposed in the hanging walls of east-vergent reverse faults (Fig. 3). While the total area of exposure is not as great in Argentina as it is in Chile (Fig. 2), the Argentine exposures are more continuous from top to bottom of the section. Moreover, the thickest deposits of the basin lie along the eastern margin of the basin in Argentina, and, therefore, the eastern sections probably provide the most complete record of the basin’s development. However, the most detailed analyses previously reported in the basin were those of Niemeyer and Muñoz (1983) and Muñoz and Niemeyer (1984), which covered only the portion of the basin exposed in the Chilean Andes and did not treat the issue of basin-forming mechanisms. Concurrent
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with the study reported here, Carpinelli-Pavisich (2000) and Radic et al. (2002) presented new structural and stratigraphic analyses from the Chilean portion of the Cura Mallín basin south of 37°S, but again, all new data were derived from a relatively shallow portion of the basin with discontinuous exposure of the stratigraphy. Consequently, this study focuses primarily on a relatively small but informative and understudied part of the eastern Cura Mallín basin, and we present here significant new sedimentologic and chronologic data that lead to a new vision of the Oligocene-Miocene–age basins of the Southern Andes. GENERAL STRATIGRAPHY OF THE BASIN The Cura Mallín basin fill consists of mostly volcaniclastic strata deposited in colluvial and alluvial systems, with minor contributions of lacustrine biogenic deposits and lava flows. Using the classification scheme of Fisher and Schmincke (1984), the volcaniclastic rocks are mostly epiclastic and reworked pyroclastic material, distinguishing material derived from volcanic rocks from that derived from unlithified pyroclastic deposits. Primary pyroclastic strata are only rarely encountered. The maximum exposed thickness of the basin is ~3 km, and the average exposed thickness of strata is ~1200 m. The age of the basin is constrained by fossil identifications and isotopic ages. Diplodon colhuapensis, freshwater bivalves of Deseadan South American Land Mammal Age (SALMA), have been identified (V. Meissinger, 1998, personal commun.) from samples taken from beds of a lacustrine sequence exposed in the Reñileuvú River valley (see Fig. 3). The Deseadan SALMA is defined as corresponding to the time range between 29 and 24.5 Ma (Flynn and Swisher, 1995). Less specific identifications of Diplodon sp. fossils have been cited, both within the study area (Uliana, 1979) and in the Chilean side of the basin (Covacevich, 1975), suggesting that lacustrine deposition was prominent in the basin. In addition to paleontologic evidence, 40Ar/ 39Ar and fissiontrack ages from the Cura Mallín Formation and a contemporaneous igneous intrusion provide additional constraints on the timing of deposition in the basin. Jordan et al. (2001) presented 40Ar/ 39Ar ages for two Cura Mallín Formation samples from the current study area (Table 1). The sample locations lie along the eastern margin of the basin between 36°45′S and 37°30′S (see Fig. 3). Additional ages from the Miocene and younger volcanic and igneous rocks in the region and from an Eocene volcanic complex (Jordan et al., 2001; this study) allow distinction of stratigraphic units and improve understanding of basin evolution. In Tables 1 and 2, we show our new ages along with those of Jordan et al. (2001) and other relevant data for samples of the Cura Mallín basin fill and the other units that occur in the study area (see Appendix 1 for 40Ar/ 39Ar age spectra). For the Cura Mallín strata, the older 40Ar/ 39Ar age (24.6 ± 1.8 Ma) is derived from a basaltic dike (sample VL-1; Jordan et al., 2001) within one of the lowermost exposed stratigraphic levels of the basin and is comparable in stratigraphic position
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Figure 2. Generalized geologic map of the Cura Mallín basin (after Jordan et al., 2001). The area shown includes the Main Cordillera, where the basin fill is exposed at the surface, as well as the Central Valley, where Cura Mallín strata occur in the subsurface (note borehole locations) and in more limited outcrop. Regional lineaments (e.g., Ramos, 1978) that cross the basin area are shown for reference.
The case for extensional tectonics in the Oligocene-Miocene Southern Andes
Figure 3. Geologic map of the study area within the eastern portion of the Cura Mallín basin in Argentina (36°45′–37°30′S, 70°35′–71°15′W). Locations of dated horizons are indicated as well. The location of a seismic line recorded in the 1980s by Argentina’s YPF company, illustrated by Jordan et al. (2001), is shown. The seismic data constrain a reverse fault, interpreted by Jordan et al. (2001) to be the northward continuation of a major inverted normal fault.
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36.738 S 71.043 W
Intrusive
unit
Granite
Rock type
Apatite
Mineral
20
Track length ±S.E. Std. Dev. No. of tracks measured (µm)
Mineral dated Age 2 error (Ma) (Ma) hornblende 3.8 0.4 hornblende 9 2 biotite 11.7 0.3 hornblende 10.8 1.6 hornblende 16.2 0.2 hornblende 22.8 0.7 hornblende 24.6 1.8 hornblende 50.3 0.6 hornblende 56.5 0.6 hornblende 56.9 1.1
Number of Central P( )2 Uranium s i d grains dated (×106 t/cm2) (×107 t/cm2) (×105 t/cm2) age (Ma) (%) content (±1 S.E.) (ppm)
TABLE 2. FISSION-TRACK AGES
Ash-flow tuff Ash-flow tuff Basalt Basaltic andesite Basaltic andesite Ash flow tuff
TABLE 1. 40Ar/39Ar ISOTOPIC AGES Longitude Stratigraphic unit Rock type (°W) 70.806 Cola de Zorro Basaltic andesite 70.918 Pichi Neuquén Basalt 70.961 Cajon Negro Ash-flow tuff
0.0876 0.623 1.241 8.9 96 60 8.2 ± 4.0 2.9 2 (60) (2134) (4304) (1.3) Zircon 20 3.31 1.49 3.2 26.3 52 721 (1097) (2464) (4600) (1.5) FT 1-1 37.226 S Cura Mallín Sandstone Apatite 25 0.0357 0.0724 12.55 28.6 92 7 12.4 ± 3.9 5.3 7 70.871 W (44) (446) (5446) (4.7) Zircon 14 3.433 1.981 2.748 15.9 <1 1118 (1461) (4215) (4046) (2.1) FT 1-3 37.225 S Cura Mallín Sandstone Apatite 20 0.03226 0.0377 12.289 51.6 40 4 n.d 70.888 W (32) (187) (5446) (11.1) FT1-4 37.226 S Cura Mallín Sandstone Apatite 20 0.03838 0.05877 12.14 38.4 35 6 n.d 70.897 W (35) (268) (5446) (8.0) FT 1-5 37.203 S Cura Mallín Sandstone Apatite 20 0.04207 0.08269 11.999 28.1 85 8 13.2 ± 0.9 2.7 36 70.937 W (35) (344) (5446) (5.2) FT 2-1 36.996 S Cura Mallín Sandstone Zircon 15 1.987 0.7957 2.7 25.1 2 457 70.968 W (412) (825) (4046) (2.4) FT2-2 37.004 S Cura Mallín Sandstone Apatite 13 0.01724 0.04483 11.868 21.1 75 5 n.d. 70.987 W (8) (104) (5446) (7.8) Notes: s—spontaneous track density, i—induced track density (reported induced track density is twice the measured density), d—track density in muscovite detector covering CN-5 (10 ppm) or CN-6 (1ppm); reported value determined from interpolation of values for detectors covering standards at the top and bottom of the reactor packages (influence of gradient correction). Number in parenthesis is the number of tracks counted for ages and influence of calibration. S.E.—one standard error. P( )2 = chi-squared probability. d = 1.551 × 10–10yr–1, g = 0.5. Zeta = 370 ± 56 for zircon; 463 ± 34 (CN5) and 4772 ± 340 (CN6) for apatite.
AL2
Sample Latitude/ number Longitude (°)
Latitude (°S) 37.209 36.932 36.997
36.937 70.715 Cajon Negro LP-4† 37.371 70.974 Cura Mallín RR-7† † 36.880 71.079 Cura Mallín VL-1 † 37.076 70.766 Serie Andesitica 97 N-volc 9 37.140 70.691 Serie Andesitica 97 N-volc 10† † 97 N-volc 13 37.275 70.706 Serie Andesitica † Sample previously published in Jordan et al. (2001).
97 N-volc 1 97 N-volc 8 RB-2
Sample number
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The case for extensional tectonics in the Oligocene-Miocene Southern Andes to the strata containing the Deseadan age (29–24.5 Ma) fossils. The composition of the dike is similar to that of lavas interbedded with volcaniclastic debris-flow deposits in this portion of the basin (Burns, 2002), so the dike is inferred to be contemporaneous with the early stages of deposition within the Cura Mallín basin. The younger age (22.8 ± 0.7 Ma) is from an ashflow deposit (sample RR-7; Jordan et al., 2001), which marks the end of significant sedimentary deposition within the study area. Therefore, the 40Ar/ 39Ar ages appear to bracket the time of deposition within the eastern Cura Mallín basin between ca. 26 and 22 Ma. Notably, ages of volcanic deposits in the Cura Mallín unit of the western outcrop belt in Chile have yielded only middle Miocene ages (Wertheim et al., 2003). Further west, however, in the Chilean Central Valley, K/Ar and apatite fission-track dating of Cura Mallín strata sampled in boreholes provides ages ranging from late Oligocene through early Miocene (Vergara et al., 1997), overlapping with those in the current study area. In general, apatite and zircon fission-track ages from sedimentary strata and one granitic body that intrudes the Cura Mallín Formation are similar to the depositional ages implied by the 40Ar/ 39Ar dates. That three out of five apatite dates and one of two zircon dates on volcaniclastic sandstones are of late Oligocene and early Miocene age implies that the volcanic sources of the detrital clastics were contemporaneous with the basin subsidence. Exceptions reflect the influence of detrital sources other than contemporaneous volcanism (FT1-3 and FT1-4) and indefinite problems with the age determination or the stratigraphic interpretation (FT1-1 zircon; see Appendix 2 for further discussion of the fission-track ages and interpretation of their significance). A late Oligocene zircon age (26.3 ± 1.5 Ma) on a granitic body (AL-2) emplaced within the lowermost exposures of the basin fill indicates that deposition within the basin must have begun by 24.8 Ma and may have begun before 27.8 Ma. A similar zircon age (25.1 ± 2.4 Ma) for a volcaniclastic deposit (FT2-1) ~200 m upsection from the pluton suggests that the pluton intruded rocks only slightly older than itself. Therefore, within the range of its error bars, the zircon age of the pluton is probably a reasonable representation of the initiation of basin formation and is consistent with the 29–24.5 Ma age assigned to fossils in the oldest known horizons. On the other hand, the apatite age for the pluton is late Miocene (8.9 ± 1.3 Ma) and reflects thermal resetting of the fission-track system, as discussed in the following. The contact of the basin fill with its basement is rarely observed, but the next-older strata in the region are Eocene-age (58–40 Ma; Rovere, 1998; Jordan et al., 2001) volcanic rocks and intercalated sedimentary deposits of the Serie Andesitica (Rapela and Llambías, 1985), also known as the Cayanta Formation (Zöllner and Amos, 1973; Rovere, 1998). These rocks are exposed immediately beyond the eastern margin of the basin in Argentina but are not observed in depositional contact with the Cura Mallín strata (Fig. 3). Whether or not this Eocene unit
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extends westward beneath the basin is unknown. Mesozoic rocks of the Neuquén Basin in the western flank of the Cordillera del Viento anticline underlie the Serie Andesitica, and the westward extent of these units beneath the Cura Mallín basin is also unknown. However, outside of the study area, in the northeastern portion of the basin, Cura Mallín strata overlie Jurassic and Cretaceous marine strata similar to those of the Neuquén Basin, as well as supposed Cretaceous plutons that intrude those strata (Muñoz and Niemeyer, 1984). Near the international boundary in the study area (Fig. 2), the basement below Cura Mallín strata is generally unknown. However, the western margin of the basin is bounded by Paleozoic rocks of the Coastal Cordillera. On the basis of gravity data, Vergara et al. (1997) proposed that the Paleozoic units extend eastward beneath the Cura Mallín strata, which were penetrated by exploration wells in the Central Valley near the town of Chillán. The upper limit of the Cura Mallín Formation is marked by a transition from primarily epiclastic deposits to the primarily volcanic deposits of the early to middle Miocene Trapa Trapa Formation defined in Chile (Niemeyer and Muñoz, 1983) and the Cajón Negro and Quebrada Honda Formations in Argentina (Pesce, 1981). Radiometric ages for these volcanic units range from ca. 22 to 11 Ma (Drake, 1976; Muñoz and Niemeyer, 1984; this study). The contact between the Cura Mallín Formation and the overlying Miocene volcanic rocks appears to be conformable where observed in the study area. However, high-resolution isotopic dating of volcanic strata in the Coya-Machalí basin (Fuentes et al., 2002) to the north has revealed both a prolonged hiatus (>8 m.y.) within an apparently continuous volcanic section and virtually continuous deposition across an angular unconformity, suggesting that unexpected time-stratigraphic relationships may be common within the Oligocene-Miocene basins. Beginning by 9 Ma, both the Cura Mallín Formation and, where they overlie it, the Miocene volcanic rocks were deformed by compressional deformation, which formed dominantly N-S–trending folds and reverse faults (Niemeyer and Muñoz, 1983; Muñoz and Niemeyer, 1984). Two new constraints are presented here for the timing of this deformation. The first is a 40Ar/ 39Ar age of 11.7 ± 0.3 Ma (Table 1, sample RB-2), which was determined for an ash-flow deposit in the footwall of one of the reverse faults that uplifts the eastern margin of the Cura Mallín basin. The fault must be younger than the 11.7 Ma rock that it cuts. The second constraint is an 8.9 ± 1.3 Ma apatite fission-track cooling age (Table 2, sample AL-2) for a granitic pluton in the hanging wall of the same reverse fault. The zircon fission-track age for this sample (26.3 ± 1.5 Ma) indicates late Oligocene emplacement, and the apatite age is best interpreted to represent cooling from a later heating event that annealed the apatite fission tracks, rather than initial cooling of the pluton. Burial beneath more than 3 km of Cura Mallín basin fill and Miocene volcanic rocks is the likely cause of the heating stage. In view of the structural position in the hanging wall of one of the basin-margin reverse faults, the cooling recorded by the apatite fission-track age was probably the result
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of exhumation of the pluton as it was uplifted on the fault. This interpretation of the fission-track age agrees well with the maximum age of deformation required by the crosscutting relationship. In addition, it places a lower age limit of 7.6 Ma on the initiation of compressional uplift. The late Miocene contractional deformation documented here is consistent with evidence for Neogene contraction in the Neuquén Basin immediately to the east of the study area (e.g., Cobbold and Rossello, 2003). Locally, late Miocene, Pliocene, and Quaternary volcanic rocks and glacial deposits unconformably overlie the deformed strata. The oldest age reported for these undeformed units is 6.1 ± 0.5 Ma (Muñoz and Niemeyer, 1984). DEPOSITIONAL FACIES Following the generalized facies associations defined by Smith and Landis (1995) for sedimentary and volcanic deposits associated with arc volcanism, two main depositional facies have been identified within the Cura Mallín basin, a volcanic apron facies and a distal volcaniclastic facies. The implied control on facies distribution is proximity to volcanic centers, where the main sediment source is syndepositional volcanism (Fig. 4). Data compiled by Smith and Landis (1995) indicate that apron facies rocks are typically found within 5–35 km of volcanic centers, while distal facies rocks are found at distances of 20–70 km from volcanic centers. In general, volcanic apron facies consist primarily of event deposits related to stripping of volcanic debris from the steep slopes of volcanoes during or soon after an eruption (Vessel and Davies, 1981; Smith, 1987; Palmer and Walton, 1990). Distal facies, on the other hand, represent terminal deposition of volcanic debris, primarily by fluvial transport, at sites that receive little direct deposition of volcanic material (Smith and Landis, 1995).
Figure 4. Idealized cross section of a composite volcano with a flanking volcanic apron, illustrating distribution of major facies as controlled by distance from the volcanic center (after Smith and Landis, 1995).
Volcanic Apron Facies Within the Cura Mallín basin, the volcanic apron facies (Fig. 5A) primarily consists of debris-flow deposits of lightly reworked pyroclastic material. The debris-flow deposits range in thickness from tens of centimeters to more than 10 m, and some of the thicker deposits can be traced laterally for more than 5 km. The rocks are poorly sorted, matrix-supported breccias with abundant angular, volcanic-lithic clasts of basaltic to dacitic composition, ranging in diameter from 0.5 cm to 20 cm. The matrix is generally tuffaceous, containing varying amounts of sand to silt-sized feldspar and volcanic glass shards with minor quartz and mafic minerals. Minor evidence for fluvial reworking of debris-flow deposits includes improved sorting and trough cross-bedding of the upper portions of some of the massive breccias. Locally, sequences of lacustrine limestones and mudstones are interbedded with the debris-flow deposits as well. These sequences locally reach thicknesses of 50 m. The limestones are micritic, laminar, dark gray to light gray, and thin to medium bedded. Mudstones are dark brown to tan with a small proportion of sand-sized lithic and feldspar grains. In addition to sedimentary rocks, lava flows and pyroclastic deposits are encountered within the apron facies. Lavas are slightly vesicular basalts and basaltic andesites, ~0.5 m to 1.0 m thick. The pyroclastic units include crystal and lithic tuffs. Crystal tuffs consist of feldspar fragments (oligoclase and andesine; Muñoz and Niemeyer, 1984), with subordinate volcanic lithics and pumice clasts in a matrix of clay and devitrified glass. Lithic tuffs include increased pumice and devitrified glass contents relative to feldspar. Volcanic lithic clasts are basaltic to dacitic in composition. Thicknesses of the tuffs range from <1 cm to ~10 cm. Aphanitic basalt and basaltic andesite dikes and sills are common at all stratigraphic levels in the volcanic apron facies. In one location, they appear to have intruded an unconsolidated debris-flow deposit that has been partially permeated and assimilated by the intrusive magma. Comparisons of the chemical compositions of the intrusive rocks to the compositions of lava flows within the basin fill and of younger volcanic rocks of the region further suggest that many of the intrusive bodies formed during deposition in the basin (Burns, 2002). However, it should be noted that, despite their predominance among the primary volcanic rocks encountered within the basin, basaltic rocks probably are not the dominant volcanic rock type produced in the Oligocene-Miocene arc. The andesitic to dacitic volcaniclastic material that comprises the majority of the volcanic apron facies has a much larger volume than do the basaltic lava flows and intrusive bodies. Therefore, the typical eruptive materials of the volcanoes contemporaneous with the Cura Mallín basin were probably andesitic to dacitic. Where exposed in the eastern subbasin, the volcanic apron facies dominates the basin fill northward from ~37°S (Fig. 2). It generally corresponds to the Río Queuco Member of the Cura Mallín Formation as described by Niemeyer and Muñoz (1983) in the western subbasin.
Figure 5. Stratigraphic columns showing representative sections of (A) the volcanic apron and (B) the distal volcaniclastic facies of the Cura Mallín Formation. The volcanic apron section is best exposed in the Buraleo River valley, and the distal volcaniclastic facies is best exposed in the Lileo River valley (Fig. 3). SALMA—South American Land Mammal Age; AFT—apatite fission-track age; ZFT—zircon fission-track age. Ages are in Ma.
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Distal Volcaniclastic Facies The distal volcaniclastic facies (Fig. 5B) primarily consists of fluvial deposits of reworked pyroclastic and epiclastic material with minor amounts of primary pyroclastic material. Medium- to coarse-grained sandstones predominate within the facies. The sandstones are feldspathic to lithic arenites and are generally immature, with >15% matrix (Muñoz and Niemeyer, 1984). Feldspar and lithic compositions are similar to those of the volcanic apron facies deposits. Bedding thickness is variable, from several centimeters to several meters, and beds are typically massive. Trough cross-bedding is common, however, in the more mature sandstones. Siltstones and mudstones are typically red to purple and form distinct sequences with little sand content. Clast compositions are similar to those of the sandstones. Lacustrine deposits are significant locally within the distal volcaniclastic facies, reaching a maximum total thickness of ~200 m. Lacustrine rocks include tan to dark gray, bioclastic, oolitic, and micritic limestones, rare black, pyrite-bearing dolostones, tan to black shales and siltstones, minor air-fall tuff, and volcaniclastic sandstones. Shales and limestones are the predominant rock types, comprising approximately equal proportions of the lacustrine sequences. Bioclastic and oolitic limestones are typically thin-bedded, but the micritic limestones reach thicknesses of ~50 cm and can be laterally extensive (up to ~10 km). Primary pyroclastic deposits in the distal volcaniclastic facies include crystal tuffs similar to those in the volcanic apron facies. However, the tuffs are uniformly thin (<2 cm) and rare. No lava flows were observed within this facies. Dikes and sills occur primarily at higher stratigraphic levels of the distal facies and are locally observed to feed directly into the extrusive rocks overlying the basin fill. Therefore, we suggest that the majority of intrusive rocks within the distal facies are related to the post–Cura Mallín volcanism (early Miocene and younger). The distal volcaniclastic facies is prominent in the eastern Andean exposures south of 37°S (Fig. 2) and generally corresponds to the Malla Malla Member of the Cura Mallín Formation as defined in the western subbasin (Niemeyer and Muñoz, 1983). EVIDENCE FOR CONTEMPORANEOUS DEPOSITION OF FACIES Based on their examination of the relatively incomplete and discontinuous stratigraphic sections available in Chile, Niemeyer and Muñoz (1983) considered the Malla Malla Member to be younger, in general, than the Río Queuco Member. They also considered the lacustrine deposits of the basin to reside within only the Malla Malla Member, our distal facies. However, the 40Ar/ 39Ar data and fission-track data for samples collected within the more complete Argentine stratigraphic sec-
tions and our observations of those sections lead to different interpretations for the eastern subbasin. The transition between the volcanic apron and distal volcanic facies is not directly observable anywhere in the Argentine sector of the basin, nor has it been cited in the Chilean sector (Niemeyer and Muñoz, 1983; Muñoz and Niemeyer, 1984). For this study, we dated samples from several stratigraphic sections of the Cura Mallín Formation in both the volcanic apron facies and the fluvial facies. The 40Ar/ 39Ar ages and fission-track data define a time range from ca. 28 to 22 Ma during which deposition occurred in the eastern part of the basin, but there is no clear indication that one facies was deposited prior to the other (Fig. 5). Rather, the two facies appear to be contemporaneous, within the resolution of the dating techniques used. Stratigraphic analysis of the basin fill provides additional evidence for coeval distal volcaniclastic and proximal volcanic apron deposition in the Cura Mallín basin. Within the distal facies, there is ample evidence for contemporaneous volcanism as a major influence on deposition. Pyroclastic air-fall deposits are found interbedded with lacustrine limestones and shales, floodplain siltstones and mudstones, and fluvial sandstones. In one location (Fig. 6A), a series of eruptions is recorded by the infilling of a lake. The lacustrine sequence is ~200 m thick and consists almost exclusively of thin- to medium-bedded lacustrine limestones and shales. Tuff layers begin to appear in the upper 10 m of the lacustrine section and increase in frequency upsection. Eventually, several sheets of slightly reworked volcanic ash cap the lacustrine sequence and initiate a sequence of fluvial sandstones. Although the volcanic eruptions did not create thick pyroclastic deposits in this part of the basin, the eventual result of the eruptions was a complete shift from lacustrine to fluvial deposition. In another case, a 300 m succession of mediumto thin-bedded, fluvial sandstones consisting of moderate- to well-sorted reworked pyroclastic material is capped by a 20-mthick, matrix-rich, tuffaceous sandstone (Fig. 6B). This body is overlain without transition by a thick sequence of oxidized siltstones and mudstones apparently deposited in a floodplain. We interpret this succession to represent channel avulsion resulting from the rapid influx of tuffaceous sediment from other portions of the basin, presumably after an eruption. Both of these cases show definite, drastic changes in the depositional system within the fluvial facies as a result of volcanic eruptions in another portion of the basin. With regard to the reported occurrence of lacustrine deposits within only the Malla Malla Member, lacustrine strata can be found within both the distal and the volcanic apron facies in Argentina. We studied one stratigraphic section (Fig. 5A) that clearly consists of the volcanic apron facies, containing thick, volcaniclastic debris-flow deposits with interbedded lava flows and pyroclastic rocks. However, lacustrine limestones are interbedded with the more typical volcanic apron units, with no evidence of a transition to a fluvial environment preceding or following the lacustrine sequence.
limestone
shale
limestone
fine
Figure 6. Detailed stratigraphic sections showing depositional variations within the distal volcaniclastic facies exposed in the Lileo River valley related to pyroclastic eruptions.
coarse
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LACUSTRINE DEPOSITS AND THEIR RELATION TO BASIN SUBSIDENCE The facies interpretations of Niemeyer and Muñoz (1983) imply that the fluvio-lacustrine Malla Malla Member represents a second phase of basin formation, during which time volcanic activity decreased and, therefore, sediment supply to the basin also decreased. Consequently, a fairly simple basin history can be inferred that requires no tectonic mechanisms for basin subsidence. Rather, initial accumulation of sediment would have resulted from local volcanism building volcanic edifices, which blocked drainages and created clastic aprons around the volcanoes. Lacustrine systems might have formed within closed intermontane basins without requiring any tectonic subsidence. A subsequent increase in volcanic activity, represented by units such as the Trapa Trapa and Cajón Negro Formations, would have then covered and preserved the Cura Mallín basin. The data presented here are inconsistent with that hypothesis and suggest an alternative basin history in which the interaction of sediment supply and variable tectonic subsidence controlled depositional variation. We have shown evidence that the volcanic apron and distal volcaniclastic facies were deposited concurrently in different parts of the eastern basin between ca. 28 and 22 Ma. Distance from the volcanic centers supplying most of the clastic material to the basin was the primary control on the distribution of these facies. However, even in locations proximal to volcanoes, which, at least occasionally, received large influxes of sediment, lacustrine systems formed and persisted. Moreover, lacustrine strata are observed at the lowest exposed stratigraphic levels of the distal volcaniclastic facies as well as in the middle and upper levels (see Fig. 5). The accumulation of the long-standing body of water implied by lacustrine deposition requires an excess of accommodation relative to sediment input (e.g., Carroll and Bohacs, 1999). Therefore, these observations indicate that, throughout the evolution of the basin, it was common for accommodation to exceed sediment supply, regardless of proximity to sediment sources. Given the correlation between lake formation and accommodation, it is believed that the distribution of lacustrine strata across the basin may provide important insight into the source of accommodation during deposition of the basin fill. The first key observation regarding the distribution of lacustrine rocks in the Cura Mallín basin is that they are typically found within the thickest stratigraphic sections in the basin and that there is a positive relationship between the total thickness of lacustrine strata within a given section and the thickness of that section. For example, between ~37°15′S and ~37°30′S, along the eastern margin of the basin in Argentina, there is a dramatic decrease in total sedimentary thickness of the basin fill and in lacustrine thickness from north to south. The Río Lileo section (Fig. 5B) of the distal volcaniclastic facies is ~3 km thick and contains more than 200 m of lacustrine limestones and shales. Along the Río Reñileuvú (Fig. 3), 20 km to the south, a section of the Cura Mallín Formation
crops out that is composed of the distal volcaniclastic facies, is only 1500 m thick, and contains <10 m of lacustrine strata. On the other hand, 30 km north of the Río Lileo section at the Río Buraleo section (Figs. 3 and 5A), the stratigraphic thickness is similar to that found in the Río Lileo section and ~50 m of lacustrine strata is present, despite the fact that this section is comprised of the volcanic apron facies. Yet, 20 km to the west, in the western part of the basin, the total stratigraphic thickness is only 1000–1200 m, and no lacustrine strata are observed (Niemeyer and Muñoz, 1983). Similarly, where deposits of the Cura Mallín basin in Chile are thickest, lacustrine strata are also found. Accordingly, in the subsurface of the Central Valley near Chillán (36.5°S), lacustrine rocks were found in a borehole that penetrates part of an estimated 4 km of Cura Mallín basin strata (Fig. 2; Vergara et al., 1997). At the same latitude, but in the Andean Cordillera, stratigraphic thicknesses of the basin fill range from 0 m in the east to ~2000 m in the west, and no lacustrine strata are reported in either the volcanic apron or the distal volcaniclastic facies (Muñoz and Niemeyer, 1984). If, as argued by Carroll and Bohacs (1999), the lacustrine rocks result from an excess of accommodation, then the correlation of lacustrine rocks to the thickest portions of the basin fill implies that accommodation was largely created by subsidence rather than by a lack of clastic sedimentary input to a volcanically dammed drainage basin. An independent source of information about the control of subsidence on the Cura Mallín basin and its lacustrine facies should be the spatial distribution of thicknesses of strata. However, a lack of knowledge of depositional timing in large parts of the basin results in uncertainty about the interpretation of thickness patterns. It is known that, between ~37°S and 38°S, the thickest stratigraphic sections of the Cura Mallín basin (~3000 m) lie within the easternmost outcrops of the basin. To the west of these lacustrine-rich outcrops, the total thickness of basin fill diminishes to 400 m, and there are no lacustrine strata (Niemeyer and Muñoz, 1983; Carpinelli-Pavisich, 2000; Radic et al., 2002). If eastern and western Cura Mallín strata span the same time interval, this thickness pattern suggests asymmetrical subsidence of an ~100-km-wide basin floor, with a concentration of subsidence along the eastern basin margin. In contrast, near 36.5°S, the Cura Mallín Formation in the Central Valley is thick and rich in lacustrine horizons, it does not crop out west of the Central Valley, and it is thinner in the Andes (Muñoz and Niemeyer, 1984). If the ages of the Central Valley subsurface strata and those cropping out in the Andes at the same latitude are the same, one could interpret this to mean that the basin tilted toward the west. This contrast between a northern sector (near 36°S) and southern sector (37–38°S) is suggestive of a pair of opposite polarity subbasins with subsidence concentrated within the western portion of the northern subbasin and within the eastern portion of the southern subbasin (Radic et al., 2002). Given that the available ages for Cura Mallín strata in the subsurface of the Central Valley (Vergara et al., 1997) are consistent with those for Argentine outcrops (Jordan et al., 2001;
The case for extensional tectonics in the Oligocene-Miocene Southern Andes this study), the proposed scenario of opposite polarity subbasins seems plausible. Even so, the lack of evidence for contemporary deposition from Cura Mallín outcrops in Chile (Wertheim et al., 2003) results in uncertainty regarding the width of the basins, leaving the possibility that the subsidence pattern may not be related to opposing polarity subbasins. Likewise, although Carpinelli-Pavisich (2000), Radic et al. (2002), and Croft et al. (2003) proposed that the Lonquimay basin south of 38°S is another west-focused subbasin, the Lonquimay strata are younger than the Argentine Cura Mallín strata (Suárez and Emparan, 1995; Carpinelli-Pavisich, 2000; Croft et al., 2003), and a simple polarity switch fails to explain the relation of the Lonquimay subbasin to the two (or more) subbasins of the Cura Mallín strata. STRUCTURAL EVIDENCE FOR TECTONICS OF BASIN SUBSIDENCE The proposed basin structure of paired, alternating-polarity subbasins is similar in form to that of basins created within some extensional regimes, most notably the East African Rift (e.g., Gibbs, 1984; Rosendahl et al., 1986). This comparison leads to the hypothesis that subsidence in the Cura Mallín basin was controlled by extensional faulting along the eastern margin of the southern subbasin and along the western margin of the northern subbasin (Fig. 7; Radic et al., 2002). To test this hypothesis, we mapped geologic structures within and around our study area in the southern Cura Mallín subbasin and attempted to understand the meaning of those structures with regard to the mechanisms of basin formation. Within the basin, the majority of the structures are N-S– trending folds and faults (Fig. 3) that are the results of the postdepositional compression that inverted the basin and the overlying middle Miocene volcanic rocks (Niemeyer and Muñoz, 1983; Muñoz and Niemeyer, 1984; this study). Few structures have been discovered that clearly precede this post–middle Miocene compressional event. Exceptions occur at the base of the Río Lileo stratigraphic section (east margin Cura Mallín exposures of Fig. 3) and immediately adjacent to the eastvergent reverse fault that serves as the eastern limit of the basin outcrop, where lacustrine strata contain two small, listric normal faults (Fig. 8). The fault planes have been rotated by post– middle Miocene compressional folding, but with the compressional effects removed, the faults trend N-S with a westward dip. Slickenlines measured on one of the fault surfaces indicate primarily dip slip, although there is a lateral component of slip as well. Apparent stratal thickening in the hanging walls of the two faults suggests syndepositional normal motion on the faults. It should be noted that the faults appear to root in detachments within the pictured sequence (Fig. 8). Therefore, they may be due to slope instability, independent of the tectonic environment. A similar relationship has been observed at a different scale in seismic-reflection data. About 20 km due north of the normalfaulted outcrop (Fig. 3), a petroleum industry seismic reflec-
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tion line shows structural relations below an Upper Miocene– Pleistocene volcanic cover. Jordan et al. (2001) interpreted these data to show that a W-dipping fault bounds the eastern margin of the Cura Mallín basin. Offset relations in the Upper Miocene volcanic units define thrust offset. Yet, below the nonreflective volcanic package, the seismic data show a relatively thick package of reflective strata in the hanging wall of the fault that is contiguous with a thinner package of reflective strata in the footwall of the fault. This stratal relationship was interpreted to indicate more than 1 km of syndepositional normal motion across the fault, much like that observed on a small scale in outcrop, and the strata contemporaneous with the normal offset were interpreted to be the Cura Mallín Formation. In sum, the seismically imaged fault was interpreted to display initial normal displacement and later reverse displacement, and thus inversion. Notably, the fault imaged in the seismic data lies directly on the northward projection of the reverse fault that bounds the Río Lileo stratigraphic section (Fig. 3). Given the evidence for normal faulting of Cura Mallín strata in the Lileo section, we propose that the Lileo fault is a continuation of the fault in the seismic data and that it is also an inverted basinbounding normal fault (Fig. 9). No other direct evidence has been found for normal-faulted basin margins in the eastern Cura Mallín basin. This paucity of evidence may result not only from the widespread cover by younger volcanics but also from erosion provoked by basin inversion, which has resulted in the removal of much of the sediment that was deposited closest to the basin margins and which was most likely to have been affected by syndepositional faulting. Nevertheless, the strata that remain do not appear to have suffered pervasive syndepositional faulting of any sort. We might also expect to see fanning of strata on a cross-basin scale if the asymmetric depositional pattern in the basin were the result of deposition in a subsiding half-graben. However, there is no location with sufficient lateral exposure of the basin stratigraphy in the east-west direction to make the appropriate observations. The compressional folding of the basin fill further complicates lateral tracing of primary bedding orientations. Despite the difficulties caused by the Miocene compressional overprinting of any syn–Cura Mallín structures, the late Miocene faults and folds may provide indirect evidence of the pre-existing basin structure. In the thin- to medium-bedded distal volcaniclastic facies of the study area, the strata are tightly folded, and deformation is strongest along the eastern basin margin, diminishing westward toward the basin interior (Fig. 9). The Eocene volcanic rocks adjacent to the eastern basin margin are nearly undeformed, despite the tight folding of the basin fill. These characteristics suggest decoupling of the OligoceneMiocene strata from their basement and a sharp, lateral transition to a more rigid body of rock adjacent to the basin, a body that may have served as a backstop to the deforming basin fill. Observations within known, inverted extensional basins show similar types of deformation of the basin fill, including detachment of the internally deformed basin fill from the underlying
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Figure 7. Schematic diagram of the Cura Mallín basin and underlying lithosphere composed of two alternating polarity half-graben basins.
basement and development of basement backstops (e.g., Gillcrist et al., 1987; Butler, 1989; de Graciansky et al., 1989). Such basement structures could potentially predate formation of the Cura Mallín basin (e.g., Mesozoic rift structures), but the stratigraphic evidence for localized subsidence proximal to the basin margins during the Oligocene-Miocene suggests that the inferred faults were active during Cura Mallín deposition. TECTONIC ENVIRONMENT OF BASIN FORMATION Other workers have proposed that strike-slip deformation played a role in the formation of neighboring Oligocene-Miocene basins (Suárez and Emparan, 1995; Spalletti and Dalla Salda, 1996; Godoy et al., 1999) and in regional Oligocene-Miocene Andean tectonics (Kay and Mpodozis, 2002). Since strike-slip deformation can also generate local extensional structures, the normal faulting cited here could have occurred within a regional strike-slip strain field. Moreover, transtension within strike-slip systems is a common mechanism of intra-arc basin subsidence
(e.g., Burkart and Self, 1985; Sarewitz and Lewis, 1991), and volcanic arcs may be ideal locations for strike-slip deformation if arc magmatism results in thermal weakening of the underlying lithosphere. Consequently, the possibility that regional strike-slip deformation generated the local extensional strain field should not be dismissed without cause. That being the case, we can assess the available data from the Cura Mallín basin and other Oligocene-Miocene basins of the Southern Andes for evidence of the tectonic regime during basin formation. Structural data for the eastern Cura Mallín basin presented here support the existence of normal faulting during basin formation and of high-angle faults along basin margins. Although limited in quantity, the kinematic data (Fig. 8) and fault offset data provide no evidence of strike-slip deformation as a significant factor in basin formation or even during the subsequent history of the Cura Mallín basin. Likewise, no direct evidence of syndepositional strike-slip faulting has been produced for any of the Oligocene-Miocene basins (e.g., Suárez and Emparan, 1995; Spalletti and Dalla Salda, 1996; Godoy et al., 1999).
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Figure 8. Field photo of small, listric normal faults within lacustrine limestones and shales of the Cura Mallín Formation that were later rotated by compressional deformation. Orientations of the fault planes and slickenlines on one of the fault surfaces are shown on a lower-hemisphere stereonet projection with the effects of compressional rotation removed by returning the hanging wall beds to horizontal. Thin lines highlight bedding surfaces; thick lines highlight faults.
Figure 9. Schematic structural cross section of the eastern margin of the Cura Mallín basin at ~37°15′S latitude. Tight folding of the OligoceneMiocene basin fill is limited to the basin margin, which consists of a hypothesized, high-angle normal fault that acted as a buttress during basin inversion. See Figure 3 for location.
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In addition to the lack of fault data suggestive of strike-slip offset, major stratigraphic characteristics of the OligoceneMiocene basins are not indicative of a strike-slip origin. The basins formed over a period of as much as 17 m.y. and typically contain no more than 3 km of sediment, giving an average accumulation rate of less than 0.2 mm/yr, much slower than typical sedimentation rates in strike-slip related basins (~3 mm/yr, Nilsen and Sylvester, 1995). Spalletti and Dalla Salda (1996) called attention to a feature common to strike-slip related basins, the tendency of depocenters to migrate parallel to the orientation of strike-slip faults through time (e.g., ChristieBlick and Biddle, 1985) and interpreted that the Ñirihuau basin and younger Lonquimay basin (Fig. 1) represent such migration due to hypothetical strike-slip faults controlling subsidence. However, the concept of lateral migration of depocenters presented by Christie-Blick and Biddle (1985) refers to the migration of a single depocenter rather than the development of multiple depocenters, so the interpretation is weakened. Moreover, at least within the Cura Mallín basin, the existing chronologic data (Vergara et al., 1997; Jordan et al., 2001; this study) are consistent with contemporaneous deposition within the deepest portions of both depocenters proposed here. Furthermore, the Argentine data imply along-strike contemporaneity of deposition within the southern depocenter. While strike-slip basins certainly may exhibit contemporaneous deposition within different depocenters, the Cura Mallín basin and the other Oligocene-Miocene basins of the region do not provide a close match to the “type” strike-slip basin described by Christie-Blick and Biddle (1985). The deformation state in the region surrounding the Oligocene-Miocene Cura Mallín basin offers little clarification of the tectonic setting. Thermal data and crosscutting relations for the Cordillera del Viento anticline (Fig. 3), near the eastern margin of the Cura Mallín basin, reveal no deformation contemporaneous with Cura Mallín basin development (Burns, 2002). Other reportedly prominent structural features adjacent to the basin margins include a group of NW- to NNW-trending lineations that have been mapped from Thematic Mapper satellite images and that are interpreted to be regionally pervasive basement faults (Fig. 2; e.g., Ramos, 1978). These lineations cut obliquely across the modern volcanic arc, apparently defining fault-bounded basement blocks (Ramos, 1978; Ramos and Folguera, 1999). Moreover, one of the lineations, which cuts across the Cura Mallín basin (the Nahueve lineament), marks a transition in the deformational style of the Neuquén Basin, separating basement-involved structures from thin-skinned deformation (Ramos, 1978; Cristallini and Allmendinger, 2000). Although a geographical relationship exists between the Cura Mallín basin and the Nahueve lineament (Fig. 2), there is no fault offset along the trace of the lineament through the basin (Niemeyer and Muñoz, 1983; Muñoz and Niemeyer, 1984; Delpino and Deza, 1995). Additionally, the Eocene rocks in the path of the lineament are generally undeformed (Zöllner and Amos, 1973; this study), and the Mesozoic strata at the
southern end of the Cordillera del Viento show no offset across the projection of the Nahueve lineation (Zöllner and Amos, 1973; Delpino and Deza, 1995). Moreover, where this and other proposed lineations cross the Chilean Central Valley, recently compiled gravity data provide little evidence of such structures in the subsurface (H.-J. Götze, 2000, personal commun.). Thus, it is unlikely that these lineaments exerted control on Cura Mallín basin subsidence. Like Charrier et al. (2002) in their review of the CoyaMachalí basin between 33°S and 36°S, we consider that the wide distribution and the thousands-of-meters thicknesses typical of Oligocene–middle Miocene basin deposits indicate widespread tectonic subsidence. This broad distribution of subsidence is typical of extensional settings, for which the primary strain induces tilting (about a horizontal axis) and basin formation. In contrast, strike-slip deformation does not change crustal thickness, and in oblique-slip (transtensional) settings, basin formation is typically a local response (Christie-Blick and Biddle, 1985). We conclude that there is no direct evidence of strikeslip control for any of the Oligocene-Miocene Southern Andes basins. The limited data that are available could be interpreted to be consistent with regional extension, as it is in the Cura Mallín basin. Folguera et al. (2002) hypothesized that the OligoceneMiocene tectonic environment of the Cura Mallín basin was characterized by homogeneous strain across the orogenic system, from forearc to retroarc zones. They envisioned that each fault within a set of faults paralleling the plate margin would accommodate both dip-slip and strike-slip offset appropriate to the net convergence vector. They attributed this homogeneous distribution of strain to the high angle of descent of the subducted plate beneath South America. Folguera et al. (2002) contrasted this condition with the tectonic environment of the Pliocene-Holocene in the Southern Andes, for which there is a high degree of partitioning of the strike-slip and dip-slip strain. In this partitioned condition, major strike-slip faults exist in the arc zone (Liquiñe-Ofqui fault), whereas dip-slip deformation dominate in the retroarc region. This recent partitioned strain field occurs above a moderately inclined subducting plate. This characterization of the partitioned tectonic environment is consistent with the prominence of strike-slip deformation within the modern Southern Andes (Cembrano et al., 1996; Diraison et al., 1998; Folguera and Ramos, 2000; Folguera et al., 2002). Calculated plate convergence vectors for the Nazca-Farallon and South American plates during the Tertiary (Pardo-Casas and Molnar, 1987; Somoza, 1998; S. Cande, 1998, personal commun.) indicate that obliquity of plate convergence was relatively low during the late Oligocene and early Miocene. Prior to and subsequent to that time, however, obliquity was notably higher, with a NE-SW orientation that would be conducive to right-lateral deformation. Thus, two quite different but perhaps complementary platescale explanations are available for widely distributed, dip-slip extension in the late Oligocene and early Miocene, when the
The case for extensional tectonics in the Oligocene-Miocene Southern Andes main depocenters of the Cura Mallín basin formed. Plate convergence directions led to subduction nearly orthogonal to the plate margin (Pardo-Casas and Molnar, 1987; Somoza, 1998; S. Cande, 1998, personal commun.), and a relatively steep angle of descent of the Nazca-Farallon plate beneath South America may have distributed the strain broadly across the orogenic belt (Folguera et al., 2002). The data suggest that E-W stretching extended the lithosphere subparallel to the plate convergence orientation, but with opposite sign. Although lacking a mechanical explanation, we note that the spatial scale of the set of Oligocene-Miocene basins, at least 1300 km parallel to the plate margin and 200 km normal to the margin, is entirely consistent with a phenomenon driven by the convergence itself. ESTIMATE OF CRUSTAL THINNING RELATED TO CURA MALLÍN BASIN FORMATION The existing structural and stratigraphic data for the Cura Mallín basin do not provide sufficient constraints for the construction of balanced geologic cross sections of the basin area and the precise estimation of amounts of crustal thinning and
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extension related to basin formation. However, a gross estimate of crustal thinning caused by the proposed extension can be determined if the problem is simplified to a simple Airy isostatic balance between generalized pre-extensional and postextensional lithospheric columns (Fig. 10). It should be noted that if lithospheric flexure is an important factor, the values presented here will be incorrect. We assume an average thickness of 1500 m for the basin fill (taken as a mean value across the ~200 km width of the basin), with an initial thickness of 35 km for the crust, based on estimates of current crustal thickness in the region of 30–40 km (Hildreth and Moorbath, 1988; Lüth et al., 2003). For simplicity, we ignore compaction of the basin fill, and we assume a uniform density of 2.4 g/cm3 for sedimentary rock. If we also assume that all subsidence was caused by sediment loading and tectonic thinning of the crust, then the tectonic subsidence can be isolated by removing the isostatic effects of the sediment load. Accordingly, Equation A (Allen and Allen, 1990) in Figure 10 provides a value for tectonic subsidence of ~0.4 km. If the tectonic subsidence is assumed to result from crustal thinning, then calculation of an Airy isostatic balance of the 0.4 km displacement
Figure 10. Airy isostatic model for crustal thinning related to Oligocene-Miocene basin formation. Tectonic subsidence (TS) equation (A) is from Allen and Allen (1990). Thickness variables take the form hx#, where the subscript indicates the component of the geologic column (i.e., s for sedimentary basin fill, c for crust, m for mantle lithosphere, and a for asthenosphere) and the numerical value indicates the time frame (1 for initial time frame and 2 for the post-loading/post-thinning time frame). Densities for the column components are indicated by the variable ρx, using the subscripts listed above. Δhx represents the change in thickness between time frames.
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provides an estimate of the amount of crustal thinning (Fig. 10). For crustal and asthenospheric densities of 2.65 g/cm3 and 3.33 g/cm3, respectively, the calculated amount of crustal thinning corresponding to 0.4 km of tectonic subsidence is ~2 km, ~5.5% of the crust’s initial thickness. The small magnitude of deformation calculated here is consistent with the scarcity of direct evidence for extensional deformation in the region.
APPENDIX 1. 40Ar/ 39Ar AGE SPECTRA Age spectra (Fig. A1) and a supporting data table (Table A1) are shown for three samples dated by the 40Ar/ 39Ar step-heating method (Mitchell, 1968) and presented here for the first time. APPENDIX 2. INTERPRETATION OF FISSION-TRACK AGES OF DETRITAL SAMPLES
CONCLUSIONS The Cura Mallín basin is one of a chain of OligoceneMiocene basins found in the Andes and Chilean Central Valley between 33°S and 45°S latitude. On the basis of geologic similarities and contemporaneity, we propose that the basins of this chain formed in response to similar tectonic mechanisms. Depositional facies, stratal thickness patterns, and geologic structures of the Cura Mallín basin are suggestive of basin formation by extensional faulting, with little contemporaneous strike-slip deformation. Previous interpretations of the majority of the other Oligocene-Miocene basins as pull-apart basins formed in strike-slip fault systems lack supporting evidence, and we suggest that the entire chain of Oligocene-Miocene sedimentary basins formed in response to extensional tectonics on the Southern Andean margin. The magnitude of the extension appears to have been small (~5.5% crustal stretching, assuming pure shear), but the effects were widespread. ACKNOWLEDGMENTS This project was funded by Repsol YPF, and by an American Association of Petroleum Geologists (AAPG) Grant-in-Aid, and a Geological Society of America (GSA) graduate student research grant to the lead author. In addition, numerous people at Repsol YPF contributed logistical support and geologic insight to this project. Special thanks are due to Repsol YPF geologists Ricardo Manoni, Ignacio Brisson, Franciso Pángaro, and Tomás Zapata for their considerable efforts to ensure our success. Suzanne Mahlburg Kay and Richard Allmendinger both contributed greatly to the work presented here, providing many key insights as well as access to their extensive knowledge of the Andes and geology. We are also indebted to HansJuergen Goetze and his colleagues at the Free University in Berlin and the Geoforschungszentrum-Potsdam for many friendly and enlightening exchanges between our two groups. Likewise, we would like to thank Ricardo Fuenzalida, Juan Pablo Radic, and Aldo Carpinelli for sharing their detailed knowledge of the Chilean side of the Cura Mallín basin and for providing a guided tour of their study area. Also, we thank Cristina Urizar, who produced the 40Ar/ 39Ar age spectra and associated data tables. Finally, technical reviews by Peter Cobbold and Brian Horton significantly improved the final presentation of the material in this paper.
Due to scarcity of primary volcanic rocks within the Cura Mallín basin fill, fission-track ages were determined for six detrital samples from the Cura Mallín Formation and for a granodiorite stock that intruded the base of the Cura Mallín Formation in an attempt to increase the density of chronostratigraphic data for the basin. Fission-track ages were determined on the basis of the density of fission tracks found within mineral grains in a sample and on the grain’s concentration of U-238, which creates the fission tracks as it decays over time. However, the fission-track ages require substantial interpretation in order to determine the significance of a reported age. The main sources of uncertainty are the possibility of provenance of mixed ages for detrital grains and the instability of fission tracks in apatite and zircon at relatively low temperatures. The closure temperatures for apatite and zircon fission-track systems are ~110 and 250 °C, respectively. As a result, all pre-existing fission tracks are erased in apatite and zircon grains exposed to temperatures above their closure temperatures, and no fission tracks accumulate until the grains have cooled below closure. While most of the fission-track ages reported for the Cura Mallín Formation are in agreement with the 40Ar/ 39Ar ages on interbedded primary volcanic deposits, several of them are inconsistent with the age of deposition of Cura Mallín strata. Two samples were derived from volcaniclastic debris-flow deposits proximal to a contemporaneous volcanic center (FT2-1 and FT2-2), and, based on their textures, they are the least likely to have undergone reworking of the original pyroclastic material. These samples provide fission-track ages (25.1 ± 2.4 Ma, 21.1 ± 7.8 Ma) that closely represent depositional ages as constrained by the 40Ar/ 39Ar data. One fluvial sandstone sample (FT1-5) also gives a fissiontrack age (28.1 ± 5.2 Ma) apparently representing volcanism contemporaneous with basin formation. However, the sampled horizon is near the top of the sedimentary section in this portion of the Cura Mallín basin (Fig. 5), which has a 40Ar/ 39Ar age of 22.8 ± 0.7 Ma. Therefore, the horizon that provided sample FT1-5 is expected to be younger than the fission-track age. Erosion of older volcanic edifices and fluvial redistribution of pyroclastic material throughout the basin probably contributed to the detrital makeup of the sampled unit. The ages reported for samples FT1-3 (51.6 ± 11.1 Ma) and FT1-4 (38.4 ± 8.0 Ma), which also derive from fluvial sandstones (Fig. 5), are older than the probable time of deposition.
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Figure A1. The 40Ar/ 39Ar age spectra for samples RB-2, 97Nvolc8, and 97Nvolc1 (Table 1). Corresponding data are given in Table A1. Shaded temperature steps indicate those steps chosen for calculation of the plateau age (average of the steps chosen, weighted by the uncertainty in the age), which is also shown. Uncertainties are stated at the 1σ level. Hbl—hornblende.
Moreover, both samples fail the chi-squared test (<1%; see Table 2), which is used to determine the probability that the dated grains came from a single age source. Therefore, we interpret these ages as representative of mixed detrital sources of variable fission-track age and, consequently, not indicative of the time of deposition. The final detrital sample, FT1-1, comes from near the base of the Cura Mallín Formation (Fig. 5) and gives an apatite fission-track age (28.6 ± 4.7 Ma) that agrees reasonably well with the 40Ar/ 39Ar defined time frame for deposition. However, the fission-track age from the relatively stable zircon system (15.9 ± 2.0 Ma) is significantly younger. The sample also fails the chi-squared test, suggesting that multiple grain-age populations contribute to the sample and that at least one of these populations is even younger than 15.9 Ma. The young zircon age should represent the minimum age of deposition for the sampled unit, but as such, it would require a stratigraphic repetition by reverse faulting, which is not supported by field observa-
tions. Due to the low zircon content of the sampled unit, this age is based on fewer grains (14) than the preferred minimum of 20 dated grains, raising some doubt about the validity of the age. The chi-squared test is also designed for a minimum sample size of 20 grains, so the test results may not be significant either. Even so, other fission-track ages reported here (FT2-1, FT2-2) are based on similar numbers of dated grains (15, 13) without causing obvious conflicts in the stratigraphic interpretation. Without additional material for dating, however, we are unwilling to accept the zircon age as a valid depositional age for the sampled horizon because of its otherwise unfounded implications for the stratigraphic interpretation. Even so, the interpretation of the lowermost 100–200 m in the Lileo stratigraphic section of the Cura Mallín Formation as a continuous succession should be accepted with caution. Due to the generally low apatite content of the samples and the apparent youth of the apatite grains that were found, none of the samples provided a sufficient number of measurable fission
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TABLE A1. 40Ar/39Ar AGE SPECTRA AND SUPPORTING DATA 36
37 38 40 40 Temp. Time at temp %40Ar* Ar/ Ar Ar/39Ar Ar/39Ar Ar/39Ar Ar*/39Ar Moles 39 Ar (°C) RB3-B biotite: Sens. (mol/mv) = 4.40E-17; J Factor = 0.004961; Weight (mg) = 2.7; Discrimination = 0.99737 ± 0.00062 26/1/7-01A 640 4.5 0.0776 0.0916 0.1486 23.6108 0.6792 4.76E-16 2.9 26/1/7-01B 700 0.6 0.0187 0.0465 0.0884 6.9178 1.4000 1.71E-16 20.3 26/1/7-01C 750 8.6 0.0128 0.0299 0.0862 5.0191 1.2461 1.17E-15 24.9 26/1/7-01D 800 8.2 0.0075 0.0238 0.0814 3.4970 1.2809 1.00E-15 36.7 26/1/7-01E 850 8.7 0.0049 0.0204 0.0775 2.7585 1.3128 1.12E-15 47.7 26/1/7-01F 900 8.2 0.0066 0.0198 0.0778 3.1886 1.2449 1.44E-15 39.1 26/1/7-01G 950 8.9 0.0066 0.0198 0.0822 3.3591 1.3933 1.20E-15 41.5 26/1/7-01H 1000 8.4 0.0075 0.0173 0.0836 3.5475 1.3285 1.44E-15 37.5 26/1/7-01I 1050 9.0 0.0073 0.0233 0.0833 3.3822 1.2311 1.32E-15 36.5 26/1/7-01J 1100 8.7 0.0042 0.0476 0.0907 2.6158 1.3763 7.93E-16 52.7 26/1/7-01K 1150 9.2 0.0014 0.0339 0.0934 1.7742 1.3658 1.08E-15 77.2 26/1/7-01L 1250 9.3 0.0006 0.1364 0.1347 1.7014 1.5280 1.11E-15 90.1 26/1/7-01M 1400 9.4 -0.0006 0.2476 0.1849 2.2370 2.4314 2.82E-16 108.9
Cum. %39Ar
Age (Ma) ± 1 s.d.
3.8 5.1 14.4 22.4 31.2 42.7 52.2 63.6 74.1 80.4 88.9 97.8 100.0
6.1 ± 2.1 12.5 ± 5.2 11.1 ± 0.8 11.4 ± 0.9 11.7 ± 0.8 11.1 ± 0.6 12.4 ± 0.7 11.9 ± 0.6 11.0 ± 0.7 12.3 ± 1.0 12.2 ± 0.8 13.6 ± 0.8 21.6 ± 3.1
8.5 26.9 9.8 41.4 26.3 9.7 2.8 15.4 20.2
56.2 74.5 81.3 86.0 88.3 91.3 95.0 96.8 100.0
9.3 ± 2.6 8.7 ± 5.2 4.1 ± 14.2 27.6 ± 20.6 57.8 ± 41.9 41.4 ± 33.6 37.1 ± 14.7 72.4 ± 16.1 55.2 ± 8.0
–225.4 70.9 72.8 44.8 187.1 52.4
0.0 8.2 90.0 96.1 97.3 100.0
5804.6 ± 8094.6 9.8 ± 6.6 10.3 ± 0.7 14.5 ± 9.1 23.9 ± 63.4 50.2 ± 36.4
26/1/20-02A 700 5.0 0.0138 0.5034 0.4535 7.6299 3.6074 8.26E-17 47.3 1.4 26/1/20-02B 800 7.4 0.0061 0.7890 0.2943 2.1859 0.4378 4.86E-16 20.1 9.9 26/1/20-02C 870 8.0 0.0107 8.9323 1.4354 5.6173 3.2151 4.08E-17 56.9 10.6 26/1/20-02D 930 8.2 0.0086 12.8050 1.4887 2.1006 0.6220 5.28E-16 29.4 19.8 26/1/20-02E 970 8.0 0.0048 13.2057 1.4561 0.7339 0.4123 3.31E-15 56.1 77.5 26/1/20-02F 1000 7.4 0.0042 12.9947 1.4011 0.8428 0.6720 5.54E-16 79.5 87.1 26/1/20-02G 1025 8.6 0.0052 16.4308 2.4058 1.2663 1.1329 2.03E-16 88.8 90.7 26/1/20-02H 1050 8.2 0.0032 16.1663 2.3108 0.6085 1.0202 2.68E-16 167.2 95.3 26/1/20-02I 1075 8.3 –0.0032 15.0516 2.0353 –0.0538 2.1620 1.36E-16 –3645.4 97.7 26/1/20-02J 1100 7.7 –0.2624 17.5555 2.4664 –38.3568 41.1446 5.14E-18 –106.0 97.8 26/1/20-02K 1125 8.7 –3.8549 58.6449 9.5088 –487.2065 683.7601 3.97E-19 –134.8 97.8 26/1/20-02L 1160 8.4 1.9530 0.0000 –0.9478 191.0118 –386.1440 –8.36E-19 –202.2 97.8 26/1/20-02M 1225 8.0 0.1307 0.3536 0.0879 39.2719 0.6789 1.13E-16 1.7 99.8 26/1/20-02N 1550 8.4 6.7931 0.0000 1.2576 1986.3850 –20.9851 1.29E-17 –1.1 100.0 Note: The irradiation of these samples was for 30.2 h and ended on 22 July 1998. These samples have the following correction factors: (39Ar/37Ar)Ca = 6.690E-04 ± 7.10E-06 (36Ar/37Ar)Ca = 2.680E-04 ± 3.60E-06 (38Ar/39Ar)K = 1.077E-02 (40Ar/39Ar)K = 4.950E-03 ± 9.0E-04
28.0 ± 10.3 3.4 ± 1.9 25.0 ± 21.2 4.9 ± 1.9 3.2 ± 0.3 5.2 ± 1.6 8.8 ± 4.3 8.0 ± 3.3 16.8 ± 6.5 295.9 ± 148.2 2483.8 ± 988.9 0.0 ± 1948.2 5.3 ± 10.1 –171.9 ± 454.4
Run ID#
39
97Nvolc8 hornblende: Sens. (mol/mv) = 4.40E-17; J Factor = 0.00494; Weight (mg) = 27.7; Discrimination = 0.99986 ± 0.00134 26/1/13-01A 26/1/13-01B 26/1/13-01C 26/1/13-01D 26/1/13-01E 26/1/13-01F 26/1/13-01G 26/1/13-01H 26/1/13-01I
750 850 910 970 1010 1040 1130 1200 1300
8.9 8.2 8.6 7.9 8.5 9.1 8.2 8.1 8.0
0.0388 0.0092 0.0148 0.0159 0.0653 0.1538 0.4865 0.1914 0.1137
1.3802 0.9706 1.5532 3.3072 10.8350 29.8474 119.3741 223.7279 137.5873
0.0955 0.1195 0.3133 0.7371 2.6128 4.3863 8.8704 5.4485 4.6988
12.4215 3.6264 4.7013 7.5343 24.8922 47.5426 137.8630 45.7121 28.2576
1.0506 0.9739 0.4590 3.1225 6.5918 4.7037 4.2046 8.2913 6.2886
5.71E-16 1.87E-16 6.92E-17 4.74E-17 2.33E-17 3.03E-17 3.75E-17 1.84E-17 3.27E-17
RB-3 hornblende: Sens. (mol/mv) = 4.40E-17; J Factor = 0.00487; Weight (mg) = 7.3; Discrimination = 0.99986 ± 0.00134 26/1/14-01A 26/1/14-01B 26/1/14-01C 26/1/14-01D 26/1/14-01E 26/1/14-01F
800 950 1030 1080 1140 1500
6.6 7.6 8.6 7.9 8.2 9.2
–24.0399 0.0051 0.0052 0.0123 0.0041 0.0330
0.0000 12.9841 13.5326 19.4931 29.8879 57.6419
–0.1126 0.7728 0.7659 2.0744 2.2708 3.4249
–2182.9857 1.5674 1.6098 3.6489 1.4369 10.6301
4920.9759 1.1175 1.1787 1.6558 2.7345 5.7961
3.12E-20 3.00E-17 3.01E-16 2.24E-17 4.54E-18 9.89E-18
97Nvolc1 hornblende: Sens. (mol/mv) = 4.40E-17; J Factor = 0.0043; Weight (mg) = 42; Discrimination = 0.99986 ± 0.00134
The case for extensional tectonics in the Oligocene-Miocene Southern Andes tracks in apatite (~100 confined tracks) for track-length analysis. However, the sample with the largest number of measured fissiontrack lengths (FT1-5, 36 track lengths) showed little evidence for track annealing. The mean track length was 13.2 ± 0.9 μm, which is only slightly below the range expected for unannealed fission tracks (Gleadow et al., 1986). Therefore, at least for this sample, annealing should not have affected the fission-track age determination. REFERENCES CITED Allen, P.A., and Allen, J.R., 1990, Basin analysis; principles and applications: Oxford, Blackwell Scientific Publications, 451 p. Burkart, B., and Self, S., 1985, Extension and rotation of crustal blocks in northern Central America and effect on the volcanic arc: Geology, v. 13, p. 22–26, doi: 10.1130/0091-7613(1985)13<22:EAROCB>2.0.CO;2. Burns, W.M., 2002, Tectonics of the Southern Andes from stratigraphic, thermochronologic, and geochemical perspectives [Ph.D. dissertation]: Ithaca, Cornell University, 218 p. Butler, R.W.H., 1989, The influence of pre-existing basin structure on thrust system evolution in the Western Alps, in Cooper, M.A., and Williams, G.D., eds., Inversion tectonics: Geological Society [London] Special Publication 44, p. 105–122. Carpinelli-Pavisich, A.A., 2000, Analisis estratigrafico, paleoambiental, estructural y modelo tectono-estratigrafico de la cuenca de Cura-Mallin, VIII y IX región, Chile–Provincia del Neuquén, Argentina [Título thesis]: Concepción, Chile, Universidad de Concepción, 158 p. Carroll, A.R., and Bohacs, K.M., 1999, Stratigraphic classification of ancient lakes; balancing tectonic and climatic controls: Geology, v. 27, no. 2, p. 99–102, doi: 10.1130/0091-7613(1999)027<0099:SCOALB>2.3.CO;2. Cembrano, J., Hervé, F., and Lavenu, A., 1996, The Liquiñe Ofqui fault zone: A long-lived intra-arc fault system in southern Chile: Tectonophysics, v. 259, p. 55–66, doi: 10.1016/0040-1951(95)00066-6. Charrier, R., Wyss, A.R., Flynn, J.J., Swisher, C.C., III, Norell, M.A., Zapatta, F., McKenna, M.C., and Novacek, M.J., 1996, New evidence for late Mesozoic–early Cenozoic evolution of the Chilean Andes in the upper Tinguiririca Valley (35 degrees S), central Chile: Journal of South American Earth Sciences, v. 9, no. 5–6, p. 393–422, doi: 10.1016/S08959811(96)00035-1. Charrier, R., Baeza, O., Elgueta, S., Flynn, J.J., Gans, P., Kay, S.M., Muñoz, 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): Journal of South American Earth Sciences, v. 15, p. 117–139, doi: 10.1016/S0895-9811(02)00009-3. Christie-Blick, N., and Biddle, K.T., 1985, Deformation and basin formation along strike-slip faults, in Biddle, K.T., and Christie-Blick, N., eds., Strike-slip deformation, basin formation, and sedimentation: Tulsa, Society of Economic Paleontologists and Mineralogists, Special Publication 37, p. 1–34. Cobbold, P.R., and Rossello, 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. Covacevich, V., 1975, Determinación paleontológica de una muestra procedente del Río Queuco, provincia del Bío Bío: Santiago, Chile, Instituto de Investigaciones Geológicas (Chile). Cristallini, E.O., and Allmendinger, R.W., 2000, Estructura de la faja plegada del Agrio, Provincia del Neuquén: Repsol YPF internal report. Croft, D.A., Radic, J.P., Zurita, E., Charrier, R., Flynn, J.J., and Wyss, A.R., 2003, A Miocene toxodontid (Mammalia Notoungulata) from the sedimentary series of the Cura-Mallín Formation, Lonquimay, Chile: Revista Geológica de Chile, v. 30, p. 3–16. deGraciansky, P.C., Dardeau, G., Lemoine, M., and Tricart, P., 1989, The inverted margin of the French Alps and foreland basin inversion, in
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Geological Society of America Special Papers Early to middle Miocene backarc magmas of the Neuquén Basin: Geochemical consequences of slab shallowing and the westward drift of South America Suzanne Mahlburg Kay and Peter Copeland Geological Society of America Special Papers 2006;407;185-213 doi: 10.1130/2006.2407(09)
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Geological Society of America Special Paper 407 2006
Early to middle Miocene backarc magmas of the Neuquén Basin: Geochemical consequences of slab shallowing and the westward drift of South America Suzanne Mahlburg Kay* Department of Earth and Atmospheric Sciences and Institute for the Study of the Continents, Snee Hall, Cornell University, Ithaca, New York 14853, USA Peter Copeland* Department of Geosciences, University of Houston, Houston, Texas 77204, USA
ABSTRACT New 40Ar/ 39Ar, major and trace element, and isotopic data for ca. 24–15 Ma backarc volcanic rocks from the Sierra de Huantraico, Sierra Negra, and Sierra de Chachahuén– La Matancilla regions (36°S–38°S) in the Neuquén Basin shed light on the early Miocene evolution of the south-central Andes. A model calling for incipient shallowing of the subducting slab under the northern Neuquén Basin and an increase in the rate of westward motion of South America relative to the underlying mantle at ca. 20 Ma can explain many regional features. Early Miocene magmatism in the Neuquén Basin began with the eruption of ca. 24–20 Ma alkali olivine basalts from monogenetic and simple polygenetic centers located up to 500 km east of the trench. Their characteristics (Ta/Hf > 0.45, εNd = +3.6–+4.2; La/Ta < 14; Ba/La < 16; 87Sr/ 86Sr = 0.7037–0.7040) indicate a backarc mantle devoid of arc-like components. These basalts erupted at a time of extension all along the margin during a period of rapid, near-normal Nazca–South America plate convergence when spreading ridges between the Pacific, Nazca, and Antarctic plates were becoming more parallel to the Chile Trench. Ridge rotation along with slab roll-back in response to slow relative motion between South America and the underlying mantle can explain why isotopically enriched magmas erupted far to the east of the trench in a generally extensional regime. Subsequently, 19–15 Ma basaltic to trachyandesitic backarc lavas with weak arc-like La/Ta (15–26), Ba/La (15–32), and Ta/Hf (0.2–4.5) ratios and a more depleted isotopic signature (εNd = +3.9–+4.7; 87Sr/ 86Sr = 0.7033–0.7037) erupted in a contractional regime. Their chemical features fit with incipient shallowing of the Nazca plate under the northern Neuquén Basin. A contractional regime that extended all along the margin can be explained by westward acceleration of South America over the underlying mantle as Nazca–South America plate convergence slowed. Keywords: Andes, plateau magmatism, Miocene, Neuquén Basin, trace elements, isotopes, plate motions, shallow subduction. *E-mails:
[email protected];
[email protected].
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, doi: 10.1130/2006.2407(09). For permission to copy, contact
[email protected]. ©2006 Geological Society of America. All rights reserved.
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INTRODUCTION Magmatism has been a common occurrence in the southern and south-central Andes since the Mesozoic rifting that led to the breakup of South America and Africa. Among these magmas are the early to mid-Miocene volcanic rocks that erupted in the Neuquén Basin (Figs. 1 and 2) after a change from slower, more oblique South America–Farallon plate convergence to more rapid, near-normal South America–Nazca plate convergence at ca. 25–24 Ma (e.g., Pardo Casas and Molnar, 1987; Somoza, 1998). This change, which is generally taken to mark the initiation of the modern Andean cycle, generally coincides with the first magmatic eruptions reaching far into the backarc of the central and eastern Neuquén Basin since the Triassic. At the same time, arc volcanism reinitiated in the Main Andean Cordillera to the west after a gap of >15 m.y. (e.g., Jordan et al., 2001). In a larger picture, the origin and evolution of these magmas are part of a late Oligocene–early Miocene story that includes the accumulation of volcanic/sedimentary sequences in forearc, intra-arc, and backarc basins before, at, and after the breakup of the Farallon plate (see Jordan et al., 2001). A number of workers have related the evolution of these sequences to the breakup of the Farallon plate and subsequent changes in the convergence regime between South America and the subducting oceanic plate (Kay et al., 1993; Muñoz et al., 2000; Jordan et al., 2001). The late Oligocene–earliest Miocene sequences have been related to an extensional setting (e.g., Muñoz et al., 2000; Jordan et al., 2001) and the younger ones to a contractional setting. The problem with a simple correlation with South American–Nazca convergence parameters is that the extensional regime has been argued to persist until ca. 20 Ma (Jordan et al., 2001) or even 13–8 Ma (Radic et al., 2002), whereas the last major change in convergence vectors is placed ca. 25 Ma (Somoza, 1998). The purpose of this paper is to put ca. 24–15 Ma backarc magmatism in the Neuquén Basin into the context of southcentral Andean evolution and to explore how a transition from a ca. 24–20 Ma extensional regime to a contractional one after 20 Ma fits with the broader tectonic picture. To do this, the spatial distribution, ages, and chemical characteristics of Neuquén Basin backarc volcanic rocks are evaluated in light of new 40Ar/ 39Ar ages, whole-rock major- and trace-element analyses, and Sr, Nd, and Pb isotopic analyses. The data are used to show that: (1) Neuquén Basin backarc volcanic rocks formerly assigned Eocene to late Miocene ages erupted from 24 to 15 Ma, (2) ca. 24–20 Ma volcanic rocks are alkali basalts with oceanicisland basalt (OIB)-like chemical signatures that erupted from monogenetic and simple polygenetic centers in an extensional setting, and (3) ca. 19–15 Ma volcanic rocks are basalts to trachyandesites-dacites with weak arc-like signatures that generally erupted from volcanic complexes in a more contractional setting. These interpretations are then used to propose that a subducted component was introduced into the backarc mantle
Figure 1. Generalized map showing distribution of latest Oligocene to early Miocene sedimentary and volcanic rocks in Chile and western Argentina between 33°S and 44°S. Early Miocene volcanic rocks in the Somuncura plateau province are concentrated west of the solid line in the outcrop region. Generalized outcrop patterns are from 1:1,000,000-scale geologic map of Chile (1980, Servicio Nacional de Geología y Minería, Santiago) and 2,500,000-scale geologic map of Argentina (1997, Servicio Geológico Minero Argentino, Buenos Aires). Extent of subsurface exposure is as in Jordan et al. (2001). Dashed box labeled Neuquén Basin shows generalized region of early Miocene volcanic rocks that are the main focus of this paper.
beneath the Neuquén Basin at ca. 20 Ma in association with the initiation of transient Miocene shallowing of the subducting Nazca plate (Kay, 2001; Kay et al., this volume, chapter 2). Finally, a regional tectonic evaluation shows that a ca. 24–20 Ma extensional regime associated with the eruption of the Neuquén Basin alkaline basalts and contemporaneous formation of forearc, intra-arc, and backarc basins at a time of rapid South America–Nazca convergence (Somoza, 1998) can be explained
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Figure 2. (A) Generalized geologic map of the Neuquén Basin region showing the distribution of Tertiary to Holocene magmatism and the Cortaderas lineament. Map is simplified from 1:2,500,000scale geologic map of Argentina (1997) and geologic maps of provinces of Mendoza (1993) and Neuquén (Delpino and Deza, 1995). Modifications in ages are in accord with radiometric dates in this volume. Names of Southern Volcanic Zone arc centers are given for reference.
by: (1) readjustments in South America–Nazca plate interactions in response to ridge rotations and ridge jumps among plates in the Pacific Ocean to the west (Tebbens and Cande, 1997; Tebbens et al., 1997), and (2) trench roll-back related to a slow margin-normal velocity for South America over the underlying mantle in the hotspot reference frame. The subsequent contractional regime after 20 Ma is shown to coincide with: (1) a time of near-parallel alignment of the Chile Trench with oceanic ridges to the west, and (2) trench advance as the South American plate began moving westward over the underlying mantle in the hotspot reference frame. The general features of the isotopic data from the early to middle Miocene Neuquén Basin volcanic rocks along with those from Oligocene to middle Miocene volcanic rocks between 33°S and 45°S latitude can be explained in the context of this model.
REGIONAL SETTING OF THE NEUQUÉN BASIN The early to middle Miocene lavas that are the focus of this study occur from 68.8°W to 70°W and from 36°S to 37.5°S, where they overlie the Mesozoic–early Tertiary sedimentary sequences that form the infill of the northern and central parts of the Neuquén Basin (Figs. 1 and 2). These lavas erupted in a backarc position relative to the ca. 24–20 Ma Cura Mallín Formation volcanic/sedimentary sequences, which formed in an intra-arc basin, and the middle Miocene arc sequences that followed (e.g., Suárez and Emparan, 1995; Jordan et al., 2001; Burns, 2002; Burns et al., this volume, chapter 8). The Cura Mallín sequences have been related to an extensional regime by Jordan et al. (2001), Radic et al. (2002), and Burns et al. (this volume, chapter 8). Within the context of the Neuquén Basin,
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Figure 2. (continued) (B) Thematic Mapper (TM) satellite image showing distribution of K-Ar and 40Ar/ 39Ar ages in early Miocene backarc volcanic rocks in the region discussed in this paper. Ages in the Sierra de Chachahuén and La Matancilla region are from González Diaz (1979) and Pérez and Condat (1996). Ages in the Sierra de Huantraico–Sierra Negra region are from Ramos and Barbieri (1988), Cobbold and Rossello (2003) and Table 1. Ages in the Cajón de Molle– El Manzano region are from Nullo et al. (2002). Dots with DR labels are localities of samples with analyses in Table 2B.
the backarc lavas erupted north of or along the Cortaderas lineament (Fig. 2; Ramos, 1978), which generally marks the southern extent of post-Oligocene backarc magmatic activity in the Neuquén Basin (Fig. 2). The significance of the Cortaderas boundary, which lies along a persistent NE-trending regional structural grain and projects into an offset in the modern Southern Volcanic Zone arc (Fig. 2), has been unclear. Kay et al. (this
volume, chapter 2) argue that this lineament marks the southern boundary of a transient Miocene shallow subduction zone. Regionally, the 24–15 Ma backarc volcanic rocks in the Neuquén Basin are part of the widespread array of Andean magmatic and sedimentary rocks from 33° to 45°S that formed in late Oligocene to early Miocene forearc, intra-arc, and backarc basins, and the later Miocene arc and backarc complexes that fol-
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Slab shallowing and the westward drift of South America lowed (see Fig. 1). To the west and south of the Neuquén Basin are ca. 29–20 Ma arc and forearc sequences, the surface and subsurface expressions of which extend from the eastern edge of the Central Valley near 36°S to the coast at 43° to 44°S (see LópezEscobar and Vergara, 1997; Muñoz et al., 2000). The presence of thick sedimentary sequences and variable intraplate to arc-like chemical affinities in associated mafic to silicic lavas has been interpreted to reflect extensive lithospheric thinning in forearc and intra-arc basins (Muñoz et al., 2000). To the east and south are the volcanic and sedimentary sequences in the Lonquimay Basin in the main Cordillera near 38°S, and the ca. 32–28 Ma arc volcanic sequences in the Ventana Formation (Rapela et al., 1988), which preceded the sedimentary sequences in the Nirihuau Basin (e.g., Cazau et al., 1987) at ~41° to 43°S. East of the Nirihuau Basin are the late Oligocene pre-plateau and voluminous plateau lavas and the early Miocene post-plateau alkaline flows, differentiates, and small-volume ultra-Na and ultra-K lavas of the Somuncura magmatic province (e.g., Corbella, 1984; Ardolino and Franchi, 1993; Kay et al., 1993, 2004). Volcanic sequences of similar age, which are not shown on Figure 1, occur to the south in the Meseta Canquel region at ~44° to 46°S (e.g., Baker et al., 1981; Marshall et al., 1986). To the north of the Cura Mallín Basin are the extensive early Miocene arc and backarc magmatic-sedimentary sequences in the Coya Machalí, which have been associated with a thin crust in an extensional to neutral tectonic regime (e.g., Charrier et al., 1996, 2002; Kurtz et al., 1997; Kay et al., 2005). Further north in the Chilean flat-slab region (not shown on Fig. 1) are the early Miocene Doña Ana arc sequences and small ca. 22 Ma Máquinas backarc olivine alkali basalt flows, which formed in a mildly extensional to neutral tectonic regime (see Kay et al., 1991; Kay and Mpodozis, 2002). DISTRIBUTION, AGES, AND CHEMISTRY OF EARLY MIOCENE BACKARC VOLCANIC ROCKS IN THE NEUQUÉN BASIN In this section, the distribution, ages, and chemistry of the 24–15 Ma volcanic sequences in the Neuquén Basin (Fig. 1) are examined. As shown on the map on Figure 2A and the Thematic Mapper satellite (TM) image in Figure 2B, these volcanic rocks occur in three regions. The first is in the general area of the Cajón del Molle and Puntilla de Huincán region in southern Mendoza, the second is in the La Matancilla–Sierra de Chachahuén region, and the third is in the Sierras de Huantraico and Sierra Negra regions (Fig. 3). Radiometric ages are listed in Table 1. Whole-rock chemical and isotopic analyses are listed in Table 2 and plotted in Figures 4–10; sample localities and descriptions are in Appendix 1. Molle Sequence in Southernmost Mendoza The early Neogene mafic volcanic sequences in the Cajón del Molle and Puntilla de Huincán region in southern Mendoza (Fig. 2) are historically important because this is the area where
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Groeber (1946 and earlier) defined the Basalto 0, Basalto Molle, and Basalto Palaoco units. These names have been widely used for the oldest mafic flows covering the Mesozoic sequences of the Neuquén Basin. More recently, Nullo et al. (2002) reexamined the volcanic stratigraphy of this region and regrouped these volcanic rocks into the Molle, Palaoco, and Puntilla de Huincán sequences of the Molle eruptive cycle. Referring to Figure 2B, the Molle sequence includes the flows in the Cajón del Molle, El Alambrado, and El Manzano regions; the Puntilla de Huincán sequence includes the flows in the Puntilla de Huincán region to the east; and the Palaoco sequence includes the flows in the Sierras de Palaoco and Barda Blanca regions to the north. Nullo et al. (2002) further correlated the Molle and Puntillo de Huincán sequences and assigned them early Miocene ages based on K-Ar dates of 17 ± 2 Ma in the Cajón del Molle and 19.4 Ma at El Manzano. The Palaoco sequence was assigned an age of less than 19.4 Ma, but greater than 13 ± 1 Ma. In the rest of this paper, only the Molle and Puntillo de Huincán sequences are addressed. They are simply referred to as the Molle sequence. Chemical analyses for five Molle and Puntilla de Huincán samples from Baldauf (1997) and Nullo et al. (2002) are reproduced in Table 2 and plotted in Figures 4–8. They include two alkali basalts (47–48% SiO2; Fig. 4) with 0.4–1.5% TiO2 and 1.3–1.4% K2O (one is fairly primitive with FeO/MgO ~ 1.1, 277 ppm Cr, and 133 ppm Ni) and two trachyandesites (54–55% SiO2; Fig. 4) with ~1% TiO2 and ~2% K2O. Their trace-element patterns in Figure 5C are consistent with a close genetic relationship between them, because: (1) the rare earth element (REE) patterns of the basalts and andesites are respectively characterized by ratios of La/Yb = 8–9 and 11–12, La/Sm = 2.3–2.6 and 5, and Sm/Yb = 2.3–2.6 and 2.4–2.6 (Fig. 6), (2) Ba/La ratios are all 19–25 (Fig. 7), (3) La/Ta ratios are all 20–24 (Fig. 7), and (4) Ta/Hf ratios all are 0.21–0.27 (Fig. 8). Along with a dacite (~64% SiO2) with 2.5% K2O, La/Yb = 24, La/Sm = 9, Sm/Yb = 2.6, Ba/La = 32, and La/Ta = 17, these samples have a number of chemical similarities to 19–15 Ma sequences in the Sierras Huantraico and Negra region discussed in the following sections (Figs. 4–8). In contrast, basaltic andesites from the Sierra de Palaoco analyzed by Nullo et al. (2002) have markedly higher La/Ta (>50) and Th/Hf (~1.5–2.5) ratios and lower Ta/Hf (~0. 7–1.0) ratios. Given that these ratios are like those of younger than 15 Ma volcanic rocks farther to the north and that their ages are considered to be younger than 19 Ma, but older than 13 Ma (see Baldauf, 1997; Nullo et al., 2002), the Palaoco sequence in this region is not considered further here. La Matancilla and Sierra de Chachahuén Regions Early Miocene alkali olivine basalt flows are widespread in the second area, which includes the La Matancilla region and the Sierra de Chachahuén of southern Mendoza (Fig. 2). The best known are the flows in the La Matancilla region that González Díaz (1979) assigned to the Palaoco Formation. The K-Ar ages
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Figure 3. Map of the Sierra de Huantraico region modified from Ramos and Barbieri (1988) showing distribution of volcanic units, radiometric ages, and principal faults and fold axis. Dots show sample localities for 40Ar/ 39Ar ages in Table 1 (light dots) and K-Ar ages (black dots) in Ramos and Barbieri (1988). General localities of 40Ar/ 39Ar ages in Cobbold and Rossello (2003) are in areas with boxed ages. Formation names are discussed in text.
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Latitude (°S) 40 39 New Ar/ Ar ages HDR18 37°15.76
TABLE 1. GEOCHRONOLOGY FOR EARLY TO MIDDLE MIOCENE VOLCANIC ROCKS Longitude Sierra de Huantraico Region Type (°W) 69°39.13
HDR25
37°35.09
69°28.57
HRD20
37°24.96
69°31.97
HDR14b
37°35.80
69°42.13
HDR4
37°33.80
69°50.00
Latitude (°S) New Ar/ Ar ages DRC21 36°36.8
Longitude (°W)
40
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Filo Morado olivine basalt. Road cut just east of Filo Morado station. Huantraico: Desfiladero Negro basaltic dike. Desfiladero Negro southeast of Sierra de Huantraico. Southern Huantraico olivine-clinopyroxene basalt flow. South of road just west of top of Bajada Ternero. Southern Huantraico hornblende-bearing andesite flow. Puesto to west of Puesto Gonzalez near Huantraico syncline axis. Cerro Villegas olivine basalt flow. Puesto Cerro de los Liebres west of Cerro Tormenta.
Age (Ma)
Groundmass
23.4 ± 0.4
Hornblende
25 ± 4
Groundmass
19.1 ± 0.8
Hornblende
19.8 ± 0.7
Groundmass
14.8 ± 1.2
Sierra de Chachahuén Region
Type
Age (Ma)
Matancilla region basalt flow South of Puesto Jumial, Hoja Matancilla (36°36.8 , 68°35.5 )
Groundmass
23.76 ± 0.08
Sierra de Chachahuén basalt flow Near Puesto Zuñiga.
Whole rock
20.4 ± 1.0
Basalt flow in southern Sierra de Huantraico. Cerro Tormenta. Cerro Sur de los Overos. Cerro Bayo de Huantraico andesite.
Whole rock Whole rock Whole rock Whole rock
36 ± 2 21 ± 2 22 ± 2 18 ± 2
39
68°35.5
K-Ar ages from Perez and Condat (1996) Ch-14 37°04.81 68° 53.99 K-Ar ages from Ramos and Barbieri (1988)
40
Ar/39Ar ages from Cobbold and Rossello (2003) Flow near Cerro Bayo de Sierra Negra. Filo Morado flow. Cerro Bayo de Sierra Negra dike. Cerro Bayo de Sierra Negra dike. Cerro Bayo de Sierra Negra dike. Cerro Bayo de Sierra Negra dike. Cerro Bayo de Sierra Negra dike.
that he reported are plotted on the TM scene in Figure 2B. They are: (1) 26 ± 2 Ma for a flow between Loma Alta and Loma Negra, (2) 24 ± 1 Ma for a flow between Mendieta and Puesto Loma Negra, (3) 23 ± 2 Ma for a flow in the Loma de las Ramadas south of Cerro Las Lajas, (4) 22 ± 5 Ma for a flow at the Loma de los Ojos de Agua north of Cerro Las Lajas, (5) 21 ± 2 Ma for a flow on the western border of Cerro Amarillo west of Puesto Peligroso, (6) 21 ± 5 Ma for a flow at Puesto Ranquil north of Cerro Las Lajas, (7) 19 ± 1, 18 ± 3, and 18 ± 6 Ma for flows in the Punta de la Barda region, and (9) 16 ± 5 Ma for a flow at Cerro El Ramblón. A new whole-rock 40Ar/ 39Ar age of 23.76 ± 0.08 Ma (DRC21) in Table 1 for the Loma Negra flow south of Puesto Jumial is in agreement with these K-Ar ages. In the Sierra de Chachahuén to the south (Figs. 2A and 2B), early Miocene alkali olivine basalt flows associated with marine-fossil–bearing sandstones and sandy conglomerates have been mapped as Chachahuén Unit 1 by Pérez and Condat (1996). They reported a whole-rock K-Ar age of 20.4 ± 1.0 Ma (Fig. 2B) for a basalt flow in this variable-thickness sequence, which reaches a maximum of 70 m near the center of the Sierra de Chachahuén. The distribution of these basalts supports an association with a NE-SW–trending fault zone, and outcropscale faults support normal fault motion (D. Ragona, 1999, personal commun.; Kay et al., this volume, chapter 10). Chemical analyses for five La Matancilla region and three Sierra de Chachahuén lava flows are presented in Table 2B. The Chachahuén samples are alkali basalts (~48% SiO2) with
22.1 ± 0.5 22.2 ± 0.5 16.1 ± 0.2 15.3 ± 0.4 18.9 ± 0.4 15.8 ± 0.1 15.2 ± 0.1
~2.2% TiO2, 1.3–1.6% K2O, FeO/MgO = 1.1–1.3, 206–330 ppm Cr, and 155–256 ppm Ni, and the La Matancilla samples are alkali basalts and hawaiites (45.6–49.5% SiO2) with ~2–3.2% TiO2, 1.4–2.1% K2O, FeO/MgO = 1.3–3.5, 190–330 ppm Cr, and 116–184 ppm Ni (see Fig. 4 and Table 2B). Among common trace-element features are overlapping REE patterns (La/Yb = 14–20, one at 29; La/Sm = 3.9–6.4; Sm/Yb = 3.7–6.4, Fig. 6B) and intraplate-like high field strength element (HFSE) signatures (La/Ta = 9–11, Fig. 7; Ta/Hf = 0.45–0.67, Fig. 8) and Ba/La (10–16, Fig. 7B) ratios. Their 87Sr/ 86Sr ratios and εNd values (0.7037–0.70740 and +4.3–+4.7, respectively; Table 3) plot in the ocean-island basalt field (Fig. 9). Sierra de Huantraico and Sierra Negra The third area where early Miocene backarc volcanic rocks occur is in the Sierra del Huantraico–Sierra Negra region (Figs. 2A and 2B). A modified version of a map of the region compiled from the 1:200,000-scale Buta Ranquil (Holmberg, 1976) and Los Chihuidos Norte (Ramos, 1981) geologic sheets by Ramos and Barbieri (1988) is shown in Figure 3. All of the Miocene and older volcanic units have been affected by the contractional deformation that produced the Huantraico syncline and related folds shown on the map. Ramos and Barbieri (1988) assigned the pre-Pliocene volcanic rocks on the map in Figure 3 Eocene to late Miocene ages. The siliceous tuffs mapped in the Carrere Formation were
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assigned an Eocene age based on a K-Ar age of 36 ± 2 Ma from an overlying lava flow near Puesto González. This flow is part of a group of widespread basaltic to mafic andesitic flows, sills, and agglomerates across the region. The oldest rocks in this group were mapped as the Lower Palaoco Formation and assigned to the Eocene based on the 36 ± 2 Ma K-Ar age mentioned already. The younger rocks in this group were designated as the Upper Palaoco Formation and assigned to the early Miocene based on K-Ar dates for flows in the La Matancilla region (González Díaz, 1979). Small subsidiary basaltic centers mapped as the Cerro Cabras Basalt were also assigned to the early Miocene based on K-Ar ages of 21 ± 2 and 22 ± 2 Ma at Cerro Tormento and Cerro Sur de los Overos. Andesitic to dacitic volcanic rocks in the Sierra de Huantraico and the Cerro Bayo de Sierra Negra were mapped as the Pichi Tril Andesite and were designated as Miocene based on a K-Ar age of 18 ± 2 Ma for a Cerro Bayo de Huantraico andesite. Finally, dikes cutting Mesozoic sedimentary sequences around the Sierra de Huantraico that were mapped as the Desfiladero Negro dikes were assigned to the late Miocene based on a 9 ± 1 Ma K-Ar age for a dike in the subsurface in the Aguada San Roque region, south of Auca Mahuída (Fig. 2; see Ardolino et al., 1996). The ages listed in Table 1 and plotted on Figure 3 show that these age assignments need to be revisited, because new 40Ar/ 39Ar ages, like four of the five K-Ar ages reported by Ramos and Barbieri (1988), indicate early to middle Miocene ages. The new 40Ar/ 39Ar dates include groundmass ages of 25 ± 4 Ma for the dike forming the Desfiladero Negro ridge, 23.4 ± 0.4 Ma for a flow in the Filo Morado ridge, 19.1 ± 0.8 Ma for a basaltic flow on the south side of the road at the Bajada Ternero, and 14.8 ± 1.2 Ma for a basalt flow near Cerro Villegas along the Cortaderas lineament. A hornblende from a mafic andesite flow near the southern axis of the Huantraico syncline yielded an age of 19.8 ± 0.7 Ma. The plateau spectra for these ages are shown in Appendix 2. Additional 40Ar/ 39Ar ages cited by Cobbold and Rossello (2003) are 22.2 ± 0.5 Ma for a basalt flow from the Filo Morado ridge, 22.1 ± 0.5 Ma for a flow and 18.9 ± 0.4 Ma for a dike just west of Cerro Bayo de Sierra Negra, and five ages ranging from 16.1 ± 0.2 to 15.2 ± 0.1 Ma for dikes surrounding the Cerro Bayo de Sierra Negra. Based on these new ages, the volcanic rocks of the Palaoco Formation, Cerro Cabras Basalt, and Desfiladero Negro dike unit of Ramos and Barbieri (1988) are reassigned to: (1) an early Miocene Filo Morado basaltic sequence that includes the Palaoco Formation flows from the Sierra Negra and the northern Sierra de Huantraico and most of the Cerro Cabras Basalt, (2) a younger early Miocene Huantraico basaltic to mafic andesitic sequence that includes the Palaoco Formation flows and Desfiladero Negro dikes from the central and southern Sierra de Huantraico, and (3) an early middle Miocene Cerro Villegas basalt. The volcanic rocks in these units along with those in the Pichi Tril Andesite and the Carrere tuffs are discussed in the following sections.
Sample SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O P2O5 Volatiles Total La Ce Nd Sm Eu Tb Yb Lu Sr Ba Cs U Th Hf Ta Sc Cr Ni Co
TABLE 2A. WHOLE ROCK CHEMISTRY: NEUQUÉN BASIN VOLCANIC ROCKS Early Miocene Molle Eruptive Cycle Cajon de Molle Puntilla Huincán Bajada Pajarito I-99 6-99 7-99 9-99 8-99 47.09 48.09 54.07 55.15 63.63 1.54 1.35 0.99 1.01 0.29 18.38 16.41 16.08 16.62 16.08 9.48 9.63 6.89 6.71 2.50 0.17 0.17 0.15 0.15 0.09 5.73 8.48 4.41 4.07 1.29 10.94 9.63 7.82 7.29 3.82 2.82 2.88 3.35 3.58 4.04 1.27 1.38 2.10 2.01 2.51 0.46 0.37 0.30 0.31 0.14 1.68 1.03 2.00 0.96 3.81 99.55 99.42 98.16 97.86 98.20 16.0 34.0 20.0 4.80 1.73 0.7 2.10 0.310 713 303 2.2 0.70 2.3 3.0 0.80 30.0 30 24 30
17.0 35.0 20.0 4.90 1.69 0.7 1.90 0.280 637 325 2.7 0.90 3.0 3.2 0.70 27.0 277 131 39
22.0 43.0 21.0 4.50 1.45 0.6 2.00 0.290 582 550 1.7 1.40 5.3 4.2 1.00 20.0 209 83 25
22.0 45.0 21.0 4.50 1.41 0.6 1.90 0.300 614 553 0.9 1.30 5.4 4.2 0.90 19.0 97 39 21
26.0 46.0 17.0 2.90 0.85 0.3 1.10 0.170 714 834 3.3 2.10 7.1 3.6 1.50 3.0 <3 <3 5
4.09 4.26 5.45 5.59 Na2O + K2O FeO/MgO 1.65 1.14 1.56 1.65 La/Sm 3.3 3.5 4.9 4.9 Sm/Yb 2.3 2.6 2.3 2.4 La/Yb 7.6 8.9 11.0 11.6 Eu/Eu* 1.02 1.02 1.02 1.02 La/Sm(chondrite) 2.03 2.11 2.97 2.97 Ba/La 18.9 19.1 25.0 25.1 Ba/Ta 379 464 550 614 La/Ta 20.0 24.3 22.0 24.4 Th/La 0.14 0.18 0.24 0.25 Ta/Hf 0.27 0.22 0.24 0.21 Th/Hf 0.77 0.94 1.26 1.29 K/Cs 4791 4242 10253 18537 Th/U 3.3 3.3 3.8 4.2 Note: Data are from Baldauf (1997) and Nullo et al. (2002).
6.55 1.94 9.0 2.6 23.6 1.02 5.46 32.1 556 17.3 0.27 0.42 1.97 6313 3.4
Filo Morado Basalt The Filo Morado basaltic sequence includes the alkali olivine basalt flows that crop out in the Sierra Negra and northern Sierra de Huantraico along the limbs of the Huantraico syncline (Fig. 3). Their ages are constrained by three 40Ar/ 39Ar ages between 23.4 ± 0.4 Ma and 22.1 ± 0.5 Ma (Table 1; Fig. 3). Analyses of three of these basalts (47–48% SiO2) in Table 2C show that their chemical characteristics overlap those of the La Matancilla–Sierra de Chachahuén flows (see Figs. 4–8). These characteristics include 2–2.3% TiO2, 1.2–1.6% K2O (Fig. 4), REE patterns characterized by La/Yb = 14–21, La/Sm = 3.8–4.5, and Sm/Yb = 3.6–4.6 (Figs. 5B and 6), and incompatible trace-element ratios characterized by La/Ta = 10–14, Ba/La = 10–12, Th/Hf = 0.5–0.8, and Ta/Hf = 0.4–0.5 (Figs. 7 and 8).
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Sample SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O P2O5 Total La Ce Nd Sm Eu Tb Yb Lu Sr Ba Cs U Th Hf Ta Sc Cr Ni Co
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TABLE 2B. WHOLE-ROCK CHEMISTRY: EARLY MIOCENE NEUQUÉN BASIN BACKARC VOLCANIC ROCKS (>20 Ma) La Matancilla region Sierra de Chachahuén Sierra de Huantraico Filo Morado DRC20B DRC20A DRC21 DRC17 DRC18 DRC4 DR19 DR58 TDR32 TDR14 HDR18 45.58 48.56 47.75 49.50 48.88 47.90 47.97 48.36 47.08 48.02 47.58 3.23 2.46 2.27 2.59 2.08 2.10 2.40 2.15 1.98 2.14 2.42 14.95 15.46 14.37 16.07 15.54 14.35 14.67 15.77 16.15 15.22 14.89 11.74 10.73 10.99 12.01 10.51 10.94 11.09 11.15 10.82 11.17 11.41 0.14 0.16 0.16 0.20 0.15 0.15 0.18 0.14 0.18 0.19 0.02 9.07 7.93 8.31 3.40 5.62 9.61 9.57 8.79 3.89 6.09 8.75 8.13 8.84 8.46 8.58 10.47 9.51 8.94 8.17 14.41 10.07 8.50 3.50 4.29 3.82 4.45 4.16 3.00 2.81 2.92 3.82 3.68 3.58 2.07 1.64 1.39 1.97 1.40 1.29 1.66 1.63 1.18 1.42 1.59 0.74 0.57 0.53 0.73 0.48 0.48 0.56 0.57 0.41 0.52 0.71 99.15 100.64 98.27 99.51 99.31 99.32 99.85 99.65 99.92 98.51 99.45 37.6 80.8 37.9 8.20 2.16 0.961 1.29 0.139 1151 358 0.2 1.2 3.5 6.2 4.1 16.3 193 184 56
26.0 52.0 28.8 6.42 1.87 0.927 1.51 0.193 870 375 0.3 0.8 2.3 4.4 2.5 20.2 364 157 51
23.1 48.5 24.5 5.88 1.73 0.862 1.39 0.164 824 283 0.3 0.8 2.1 4.1 2.3 19.1 327 216 51
32.4 70.7 34.0 7.96 2.22 1.047 1.62 0.194 1014 364 0.2 1.0 3.1 5.9 3.1 18.8 275 167 44
21.7 47.1 25.3 5.63 1.62 0.843 1.52 0.200 816 314 0.2 0.8 2.3 4.2 1.9 20.4 216 116 45
19.8 44.2 17.8 5.04 1.73 0.774 1.45 0.178 831 254 0.2 0.5 1.8 3.8 1.9 19.9 330 256 57
25.7 54.0 18.0 6.00 1.85 0.844 1.38 0.207 1061 413 0.2 1.5 2.5 4.6 2.6 19.2 308 231 55
5.57 5.93 5.21 6.42 5.56 4.29 4.47 Na2O + K2O FeO/MgO 1.29 1.35 1.32 3.53 1.87 1.14 1.16 La/Sm 4.6 4.0 3.9 4.1 3.9 3.9 4.3 Sm/Yb 6.4 4.3 4.2 4.9 3.7 3.5 4.3 La/Yb 29.2 17.3 16.6 20.0 14.3 13.6 18.5 Eu/Eu* 0.91 0.94 0.94 0.93 0.91 1.08 1.00 La/Sm(chondrite) 2.79 2.46 2.39 2.48 2.35 2.39 2.60 Ba/La 9.5 14.5 12.2 11.2 14.5 12.8 16.1 Ba/Ta 87 148 126 119 163 135 161 La/Ta 9.1 10.3 10.3 10.6 11.3 10.6 10.0 Th/La 0.09 0.09 0.09 0.09 0.10 0.09 0.10 Ta/Hf 0.67 0.57 0.55 0.52 0.46 0.49 0.56 Th/Hf 0.56 0.52 0.52 0.53 0.54 0.47 0.55 K/Cs 72,469 40,761 40,807 69,210 64,044 55,795 57,491 Th/U 3.0 2.9 2.7 3.0 3.0 3.5 1.7 37°05.44 Latitude (°S) 36°31 36°36.8 36°50.2 36°49.62 37°3.72 36°31 Longitude (°W) 68°32 68°35.5 69°04.07 69°1.03 68°54,72 68°51.74 68°32 Elevation (m) 1226 1396 ± 40 Note: Sample localities are in Appendix 2. Analytical methods are discussed in Appendix 3.
Also included in the Filo Morado sequence are smallvolume clinopyroxene-bearing basaltic flows and necks from the Cerro Cabras Basalt of Ramos and Barbieri (1988), which occur south and southeast of the Sierra de Huantraico. An early Miocene age for this group is based on the K-Ar ages of 21 ± 2 Ma from Cerro Tormenta and 22 ± 2 Ma from Cerro Sur de los Overos in Ramos and Barbieri (1988; Fig. 3). An analysis of an alkali basalt (48% SiO2) from Cerro Cabras differs from the Filo Morado basalts to the north in having somewhat lower La/Yb (9), La/Sm (3), Sm/Yb (3.1), Ta/Hf (0.38), and Th/Hf (0.40) ratios and a higher Ba/La (19) ratio (Table 2B; Figs. 4–8). Huantraico Basalt and Mafic Andesite Group The Huantraico sequence includes the basaltic to basaltic andesitic flows, sills, and dikes in the southern and central
HDR17 48.38 2.23 15.01 12.14 0.17 8.24 8.75 4.30 0.94 0.37 100.53
25.7 53.5 24.0 5.96 1.86 0.823 1.53 0.187 1684 326 0.2 0.9 2.9 4.6 2.4 20.5 206 155 54
22.3 47.5 28.1 5.83 1.76 0.780 1.63 0.216 787 268 0.2 0.6 2.9 4.0 1.6 21.9 209 88 48
22.6 48.3 26.0 5.82 1.69 0.785 1.43 0.184 656 270 0.4 0.6 2.4 4.5 2.3 18.8 241 199 53
35.1 73.0 40.0 7.74 2.20 0.967 1.70 0.237 855 346 0.6 0.9 4.1 5.3 2.8 20.3 260 160 51
17.1 41.4 23.8 5.80 1.87 0.913 1.84 0.237 929 322 0.4 0.6 1.7 4.2 1.6 24.9 262 137 53
4.55 1.27 4.3 3.9 16.8 1.02 2.63 12.7 133 10.5 0.11 0.53 0.62 64,064 3.2 37°5.74 68°56.61
5.00 2.78 3.8 3.6 13.7 1.00 2.33 12.0 165 13.7 0.13 0.40 0.72 42,780 5.1 37°12.61 70°8.83 2158 ± 30
5.10 1.84 3.9 4.1 15.8 0.95 2.36 11.9 118 9.9 0.11 0.51 0.53 32,145 3.9 37°10.50 69°38.7
5.18 1.30 4.5 4.6 20.6 0.96 2.76 9.9 123 12.4 0.12 0.53 0.78 23,599 4.7 37°15.76 69°39.13 917 ± 25
5.24 1.47 2.9 3.1 9.3 1.01 1.79 18.8 203 10.8 0.10 0.38 0.40 18,287 2.8 37°31.32 69°52.22
Sierra de Huantraico region, as well as the Desfiladero Negro dike unit of Ramos and Barbieri (1988). The 40Ar/ 39Ar ages, which range from 25 ± 4 Ma to 18.9 ± 0.4 Ma, along with the K-Ar ages show that most, if not all, of these rocks are early Miocene in age. The question is the flow dated at 36 ± 2 Ma by Ramos and Barbieri (1988). Importantly, this flow is from the same region as the Huantraico flow dated at 19.7 ± 0.7 Ma (Table 1; Fig. 3). Possibly the flow dated at 36 ± 2 Ma comes from a small center on the northern edge of the Cortaderas lineament. If so, this is the only flow that has yielded a late Eocene to Oligocene radiometric age in the middle to far backarc of the Neuquén Basin. The most primitive basalts in the Huantraico sequence occur as flows and dikes in the central and eastern part of the Sierra de Huantraico and as dikes cutting nearby Cretaceous sedimentary
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Sample SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O P2O5 Total La Ce Nd Sm Eu Tb Yb Lu Sr Ba Cs U Th Hf Ta Sc Cr Ni Co
S.M. Kay and P. Copeland TABLE 2C. WHOLE-ROCK CHEMISTRY: EARLY MIOCENE NEUQUÉN BASIN BACKARC BASALTS (<20 Ma) Sierra de Huantraico basaltic flows Sierra de Huantraico basaltic dikes HDR6A HDR19 HDR21 HDR20 HDR8 HDR7 HDR2 HDR6C HDR24 HDR25 HRD11 49.01 48.01 48.78 47.17 49.08 49.59 47.38 48.90 47.76 48.63 49.17 1.65 1.54 1.61 1.69 1.55 1.20 1.59 1.69 1.66 1.50 1.56 17.86 18.66 17.88 19.22 18.28 16.96 16.95 19.43 17.20 17.98 18.20 11.45 10.49 9.79 11.17 10.28 9.91 10.49 10.11 10.12 10.92 9.61 0.17 0.17 0.19 0.19 0.15 0.16 0.17 0.02 0.18 0.18 0.21 3.69 3.94 4.49 4.83 3.93 7.54 9.12 4.11 6.45 5.38 5.54 11.74 11.74 11.32 11.57 11.27 11.52 11.10 10.09 9.97 10.81 8.76 3.09 3.23 3.51 3.01 3.37 2.69 2.59 3.27 3.76 3.34 4.33 1.02 1.15 1.46 0.77 1.03 0.64 1.06 1.53 1.25 1.23 1.65 0.28 0.30 0.50 0.27 0.28 0.22 0.28 0.33 0.52 0.32 0.40 99.91 99.23 99.53 99.90 99.21 100.45 100.72 99.47 98.86 100.28 99.44 13.1 30.3 17.1 4.67 1.50 0.743 1.91 0.257 695 254 0.4 0.4 1.8 2.7 0.8 34.8 104 43 36
15.0 35.8 20.9 5.12 1.45 0.818 2.14 0.334 834 192 0.3 0.4 1.7 2.4 0.7 22.4 28 23 37
19.9 44.8 21.9 5.56 1.69 0.819 2.02 0.289 785 350 0.7 0.8 2.5 3.3 1.3 29.1 118 46 35
11.7 27.3 14.5 4.54 1.43 0.818 2.30 0.298 678 128 0.9 0.4 1.6 2.3 0.7 25.0 16 22 40
13.7 31.6 17.6 4.55 1.50 0.777 2.22 0.292 710 212 0.2 0.6 1.5 2.7 0.8 29.4 65 30 34
10.4 26.9 15.9 3.88 1.26 0.677 1.91 0.272 592 154 0.2 0.3 1.2 2.1 0.5 39.5 204 57 43
12.8 30.2 16.3 4.22 1.36 0.698 1.68 0.235 692 288 0.5 0.6 2.1 2.7 0.8 37.4 251 92 48
18.0 41.1 22.7 5.51 1.64 0.860 2.34 0.364 808 300 1.1 1.1 3.0 3.1 1.2 18.3 9 13 32
21.6 46.1 24.1 5.68 1.82 0.862 2.00 0.265 926 967 0.8 0.8 2.5 3.4 1.4 24.8 115 52 38
4.10 4.38 4.97 3.78 4.40 3.33 3.65 4.80 5.01 Na2O + K2O FeO/MgO 3.10 2.66 2.18 2.31 2.62 1.31 1.15 2.46 1.57 La/Sm 2.8 2.9 3.6 2.6 3.0 2.7 3.0 3.3 3.8 Sm/Yb 2.4 2.4 2.8 2.0 2.1 2.0 2.5 2.4 2.8 La/Yb 6.9 7.0 9.8 5.1 6.2 5.5 7.6 7.7 10.8 Eu/Eu* 1.00 0.88 0.97 0.94 1.00 0.98 0.99 0.93 1.01 La/Sm(chondrite) 1.71 1.78 2.17 1.56 1.84 1.64 1.85 1.99 2.32 Ba/La 19.3 12.8 17.6 11.0 15.4 14.7 22.4 16.6 44.7 Ba/Ta 302 266 269 193 265 310 340 247 669 La/Ta 15.6 20.8 15.3 17.6 17.2 21.0 15.2 14.9 15.0 Th/La 0.14 0.11 0.13 0.14 0.11 0.12 0.17 0.17 0.11 Ta/Hf 0.31 0.30 0.40 0.29 0.29 0.24 0.31 0.39 0.43 Th/Hf 0.67 0.70 0.77 0.69 0.54 0.58 0.79 0.96 0.74 K/Cs 18,838 27,193 18,576 6932 39,424 30,313 16,153 11,191 12,785 Th/U 4.8 4.7 3.1 4.3 2.4 4.0 3.5 2.8 3.2 Latitude (°S) 37°31 37°23 37°31.24 37°31.33 37°31.24 37°23.98 37°24.31 36°59.52 37°31.0 Longitude (°W) 69°38.02 69°30.93 69°37.70 68°53.29 69°33 69°33 69°23 69°38.02 68°26.82 Elevation (m) 1283 ± 22 1239 ± 30 1222 ± 27 1450 ± 30 1286 ± 27 1283 ± 22 1842 ± 35 Note: Sample localities are in Appendix 2. Analytical methods are discussed in Appendix 3.
strata. Many are characterized by large clinopyroxene phenocrysts. Compositionally, they are tholeiitic to alkalic (47–50% SiO2; Table 2C) basalts that differ from the Filo Morado basalts in having lower TiO2 (1.2–1.7) and K2O (0.6–1.6) contents; flatter REE patterns (La/Yb = 5–11; Fig. 6A) and less heavy (H) REE depletion (Sm/Yb = 2–2.8, Fig. 6B); higher La/Ta ratios (15–21; Fig. 7A); lower Ta/Hf (most 0.3–0.42) and Th/Hf (0.5–1.1) ratios (Fig. 8A), and lower Th concentrations (Fig. 8B). They also tend to have higher Ba/La (up to 20) ratios (Fig. 7B). Overall, the compositions of the Huantraico basalts are intermediate between the Filo Morado, La Matancilla, and Sierra de Chachahuén basalts and the Molle basalts (compare in Tables 2A to 2D and Fig. 5). The Huantraico sequence also includes evolved clinopyroxene-bearing basalt and clinopyroxene-hornblende–bearing
15.5 32.4 20.5 4.90 1.51 0.731 2.13 0.284 654 242 0.7 0.6 2.2 2.6 0.9 27.4 39 34 41
20.7 44.4 24.2 5.73 1.68 0.817 2.14 0.279 896 418 2.2 1.2 3.4 3.1 1.3 25.1 42 23 34
4.57 2.03 3.2 2.3 7.3 0.98 1.92 15.7 270 17.3 0.14 0.34 0.85 14,171 3.5 37°35.09 69°28.57 781 ± 40
5.98 1.73 3.6 2.7 9.7 0.95 2.20 20.2 331 16.4 0.16 0.40 1.07 6182 2.8 36°59.52 68°53.29 1450 ± 30
HDR23 49.52 1.47 18.66 10.75 0.19 3.98 10.33 3.74 1.40 0.30 100.34
HDR13 49.00 1.47 19.08 10.24 0.20 4.42 9.10 4.11 1.35 0.33 99.30
16.4 38.2 18.4 4.82 1.50 0.774 2.43 0.324 872 292 1.2 0.9 2.6 2.9 0.9 24.0 18 <3 36
16.3 37.1 20.8 5.01 1.46 0.743 2.32 0.313 712 311 2.0 0.9 2.7 2.8 0.9 23.1 11 21 35
5.14 5.46 2.70 2.32 3.4 3.2 2.0 2.2 6.8 7.0 0.96 0.92 2.08 1.97 17.8 19.1 314 339 17.6 17.7 0.16 0.17 0.32 0.33 0.87 0.97 9783 5643 2.7 2.9 37°38.18 37°41.8 69°35.83 69°34 873 ± 33
basaltic andesite flows and dikes (51.3–55% SiO2). These rocks crop out in the prominent cliffs at the southern end of the Sierra de Huantraico and as small eroded eruptive centers on the plateau. Their chemical characteristics (Table 2D) are consistent with a genetic relation with the basalts. Overall, they plot near or in the trachyandesitic field on the alkali-silica diagram (Fig. 4A), have 1.0–1.7% TiO2 and 1.4–2.9% K2O, show a tholeiitic trend on the Miyashiro FeO/MgO versus SiO2 diagram (Fig. 4B), have similar or only slightly steeper REE patterns (La/Yb = 7–12; La/Sm = 3.4–4.6; Sm/Yb = 2.1–2.7; Figs. 6A and 6B), and overlapping La/Ta (16–21), Ba/La (most 16–21), Th/Hf (0.7–1.3), and Ta/Hf (0.3–0.4) ratios (Figs. 7 and 8). Higher La/Sm and Sm/Yb ratios in the basaltic andesites (Fig. 6B) fit with the removal of amphibole pheno-
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TABLE 2D. WHOLE-ROCK CHEMISTRY: EARLY MIOCENE NEUQUÉN BASIN BACKARC VOLCANIC ROCKS (<20 Ma). Sierra de Huantraico region—basaltic andesitic flows and dikes Pichi Tril group “Carrere” Sample SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O P2O5 Total La Ce Nd Sm Eu Tb Yb Lu Sr Ba Cs U Th Hf Ta Sc Cr Ni Co
HDR9-1 53.19 1.60 19.50 6.85 0.24 2.47 8.59 4.07 2.94 0.50 99.97
HRD6B 51.87 1.59 18.49 9.97 0.17 2.81 8.95 4.04 1.43 0.27 99.59
HDR22 51.26 1.47 19.52 8.18 0.19 3.38 8.75 4.14 1.74 0.40 99.04
HDR1 52.27 1.26 21.66 9.53 0.19 0.68 7.61 4.27 1.77 0.55 99.80
HDR14a 52.57 1.17 18.29 8.39 0.15 3.53 8.53 4.05 1.64 0.39 98.71
HDR15 54.40 0.99 17.80 8.21 0.16 2.85 8.74 4.03 2.21 0.33 99.74
HDR14B 55.14 0.99 19.00 7.07 0.15 2.47 7.36 4.57 2.03 0.44 99.17
HDR16 55.16 1.04 18.44 7.46 0.14 2.17 7.43 4.17 2.32 0.31 98.64
HDR6d 59.04 0.71 19.06 5.80 0.05 1.57 4.94 4.77 3.53 0.29 99.76
HDR10 61.23 0.82 17.04 5.24 0.13 1.92 5.56 4.65 3.13 0.25 99.96
HDR3A 58.84 0.96 19.97 4.26 0.12 0.63 7.73 4.78 1.99 0.69 99.97
HDR3B 67.13 0.43 16.38 3.18 0.12 0.82 4.54 4.55 2.85 0.12 100.12
Cerro Villegas HDR4 47.87 3.37 15.05 11.32 0.15 7.04 8.05 3.48 2.91 1.00 100.24
30.7 62.1 29.1 6.73 1.90 0.934 2.53 0.371 834 553 1.8 1.9 5.1 4.0 1.6 18.6 31 24 28
13.9 29.7 15.6 4.08 1.26 0.660 1.92 0.261 578 218 0.2 0.6 1.8 2.7 0.8 16.1 8 5 22
22.3 48.7 30.2 6.21 1.85 0.892 2.37 0.320 862 412 0.4 1.1 2.9 3.7 1.4 12.2 4 4 21
26.2 56.4 30.4 7.20 2.23 1.123 3.25 0.425 916 946 0.1 1.3 2.7 4.4 1.4 15.4 14 11 23
23.8 52.4 25.2 6.10 1.78 0.855 2.29 0.325 780 346 0.3 1.0 3.1 3.9 1.2 14.8 16 13 25
25.0 51.8 25.3 5.61 1.49 0.736 2.48 0.357 746 526 0.6 2.1 5.4 4.6 1.6 17.9 51 25 24
28.0 61.1 31.1 6.37 1.69 0.767 2.40 0.324 802 462 0.7 1.6 4.1 4.3 1.4 9.4 6 7 18
25.0 50.0 21.8 5.40 1.55 0.764 2.43 0.332 683 458 0.8 1.6 4.9 4.2 1.5 12.2 6 9 19
31.6 65.7 26.8 5.12 1.30 0.589 2.18 0.317 603 526 0.5 1.9 6.1 5.1 2.1 6.5 9 7 8
31.4 66.8 27.8 5.41 1.28 0.650 2.16 0.331 658 437 1.2 3.3 7.9 4.7 2.2 7.4 9 7 12
39.6 82.3 41.3 7.43 1.87 0.834 2.62 0.341 1008 705 0.3 2.1 6.0 6.0 2.4 5.2 2 4 7
30.7 52.2 12.7 3.34 0.89 0.410 1.64 0.191 519 495 0.7 2.2 6.4 4.2 1.9 3.3 6 5 5
41.6 93.2 45.6 9.94 3.04 1.200 1.31 0.145 1363 430 0.5 1.6 4.2 6.9 4.3 14.8 179 137 48
7.41 3.89 9.2 2.0 18.7 0.90 5.58 16.1 263 16.3 0.21 0.45 1.52 34,247 2.9 37°23.2′ 69°21.74′
6.39 1.61 4.2 7.6 31.7 1.04 2.54 10.4 101 9.8 0.10 0.61 0.61 46,226 2.6 37°33.8′ 69°50′
7.02 5.46 5.88 6.04 5.69 6.25 6.60 6.49 8.29 7.78 6.78 Na2O+K2O FeO/MgO 2.77 3.55 2.42 14.10 2.37 2.88 2.87 3.44 3.70 2.73 6.77 La/Sm 4.6 3.4 3.6 3.6 3.9 4.5 4.4 4.6 6.2 5.8 5.3 Sm/Yb 2.7 2.1 2.6 2.2 2.7 2.3 2.6 2.2 2.3 2.5 2.8 La/Yb 12.1 7.2 9.4 8.1 10.4 10.1 11.7 10.3 14.5 14.5 15.1 Eu/Eu* 0.92 0.96 0.96 0.97 0.95 0.88 0.91 0.93 0.88 0.81 0.81 La/Sm(chondrite) 2.77 2.07 2.19 2.21 2.37 2.71 2.68 2.81 3.75 3.54 3.24 Ba/La 18.0 15.7 18.4 36.1 14.5 21.1 16.5 18.4 16.7 13.9 17.8 Ba/Ta 337 286 303 688 286 326 341 297 252 199 292 La/Ta 18.7 18.2 16.4 19.1 19.7 15.5 20.7 16.2 15.1 14.3 16.4 Th/La 0.17 0.13 0.13 0.10 0.13 0.22 0.15 0.20 0.19 0.25 0.15 Ta/Hf 0.41 0.29 0.37 0.31 0.31 0.35 0.31 0.37 0.41 0.47 0.40 Th/Hf 1.28 0.69 0.78 0.62 0.81 1.18 0.95 1.17 1.20 1.69 0.99 K/Cs 13,938 52,893 35,402 129,413 53,622 28,417 25,152 24,701 54,167 22,520 52,160 Th/U 2.7 3.2 2.6 2.1 3.0 2.6 2.5 3.0 3.1 2.4 2.9 Latitude (°S) 37°31′ 37°31.24′ 37°30.8′ 37°23′ 37°35.8′ 37°32.2′ 37°35.8′ 37°28.76′ 37°31.24′ 37°31.0′ 37°31.32′ Longitude (°W) 69°33′ 69°38.02′ 69°36.72′ 69°24.8′ 69°42.13′ 69°42.47′ 69°42.13′ 69°40.88′ 69°38.02′ 69°33′ 69°52.22′ Elevation (m) 1286 ± 27 1283 ± 22 1286 ± 27 1650 ± 50 1283 ± 22 Note: Sample localities are in Appendix 2. Analytical methods are discussed in Appendix 3.
crysts. A temporal as well as chemical relation with the basalts is consistent with similar ages for a hornblende andesite flow (HRD14b, 19.8 ± 0.7 Ma) near the base of the southern cliff and a basaltic flow (HRD20, 19.1 ± 0.7 Ma) on the eastern side of the plateau. As shown in Figure 9 and Table 3, the Huantraico basalts are isotopically distinct from the Filo Morado basalts in having lower 87Sr/ 86Sr (0.7032–0.7033) ratios and higher εNd values (+4.3–+4.7). Within the Huantraico sequence, the basaltic samples have lower 87Sr/ 86Sr and 206Pb/ 204Pb ratios and higher εNd values than a basaltic andesite sample (87Sr/ 86Sr = 0.70365; εNd = +3.9). Pb isotopic ratios of the Huantraico samples fall in the intraplate basalt field (Fig. 10).
Pichi Tril Sequence The Pichi Tril sequence includes the andesitic and dacitic units in the Sierras de Huantraico and Negra region (Fig. 3) that Ramos and Barbieri (1988) assigned an early Miocene age based on the K-Ar date of 18 ± 2 Ma at the Cerro Bayo de Huantraico. Within error, this age overlaps those of the youngest Huantraico flows (19.1 ± 0.8 Ma; Table 1) and the oldest dikes cutting the Cerro Bayo de Sierra Negra (16.1 ± 0.2 Ma; Cobbold and Rossello, 2003). Analyzed samples in this group are from the Cerros Bayo de Huantraico and Bayo de Sierra Negra (Table 2D). All are partially altered. Two are trachyandesites with 58–61% SiO2, and a third, a dacite with 67% SiO2 (Fig. 4). Their immobile
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Figure 4. (A) K 2O+Na 2O (wt%) versus SiO2 (wt%) in early Miocene Neuquén backarc volcanic rocks plotted on classification diagram for alkaline igneous rocks. Data for Molle samples (open diamonds) are from Baldauf (1997) and Nullo et al. (2002). Data for Sierra de Huantraico–Sierra Negra region (diamonds) and La Matancilla (dark squares)–Sierra de Chachahuén (open squares) region samples are from Tables 2A to 2D. (B) FeO/MgO (wt%) versus SiO2 (wt%) for same samples as in part A plotted relative to the tholeiitic and calc-alkaline fields of Miyashiro (1974).
element compositions resemble those of evolved magmas commonly associated with mafic flows like those in the Huantraico sequence (Fig. 5C). Features in common with the mafic rocks include similar heavy REE slopes (Sm/Yb = 2–2.8; Fig. 6), and incompatible element ratios (La/Ta = 14–16; Ba/La = 14–18; Ta/Hf = 0.41–0.47; Figs. 7 and 8). Trends to higher La/Yb, La/Sm, and Ta/Hf (0.41–0.47) ratios resemble those in trachyandesitic to dacitic magmas in which residual hornblende and
accessory phases have preferentially removed middle REEs, Hf, and Ta. Their overall features have similarities to a Puntilla Huincán dacite in the Molle sequence (Table 2A; Figs. 4–8). Carrere Formation Tuffs The least understood volcanic rocks in the Sierra de Huantraico region are the Carrere Formation tuffs. Ramos and Barbieri (1988) assigned an Eocene age to the Carrere Forma-
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Figure 5 (on this and following two pages). (A–C) Extended trace-element plots for early Miocene backarc volcanic rocks from the Neuquén Basin. Concentrations (ppm) for primitive mantle normalization are from Sun and McDonough (1989): Cs (0.032), Ba (6.989), Th (0.085), U (0.021), K (250), Ta (0.041), La (0.687), Ce (1.775), Sr (21.1), Nd (1.354), Sm (0.444), Hf (0.309), Eu (0.168), Tb (0.108), Yb (0.493), and Lu (0.074). Concentrations for chondrite normalization for rare earth elements (REEs), Sr, Th, and Ta plot are from the Leedey chondrite: La (0.378), Ce (0.976), Nd (0.716), Sm (0.23), Eu (0.0866), Tb (0.0589), Yb (0.249), Lu (0.0387), Th (0.05), Ta (0.02), and Sr (116). All data are from Table 2, except where indicated on plot.
tion based on the supposition that the overlying mafic flows were Eocene in age and from mammal fossils. Both criteria have problems. First, as discussed already, most (if not all) of the mafic flows have early Miocene ages. Second, the fossils are similar to those in the Collón Cura Formation in the Nirihuau Basin (Pascual et al., 1978; Fig. 1), which is now known
to have a Miocene age (Cazau et al., 1987). An early Miocene age that makes the Carrere Formation temporally correlative with the Pichi Tril sequence is consistent with a nonerosional contact between the Carrere tuffs and overlying mafic lava flows that likely belong to the Huantraico sequence. Although no samples from tuffs mapped as Carrere Formation by Ramos
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Figure 5 (continued).
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Figure 5 (continued).
(1981) were analyzed, a clast from a tuff below the Huantraico flows in the Bajada del Toro (Fig. 3) is chemically like the Pichi Tril Andesite from Cerro Bayo de Huantraico. Overall, an early Miocene age for the Carrere Formation tuffs seems most logical. Nevertheless, it is still possible that these tuffs are composed of pyroclastic material derived from Eocene centers on the east side of the Cordillera del Viento to the west (Fig. 2; Llambías and Rapela, 1989; Cobbold and Rossello, 2003). Cerro Villegas Basalt A new 40Ar/ 39Ar age of 14.8 ± 1.5 Ma (Table 1) for a basalt near Cerro Villegas along the Cortaderas lineament (Fig. 3) shows that not all flows assigned to the Cerro Cabras Basalt by Ramos and Barbieri (1988) have early Miocene ages. The chemistry of the Cerro Villegas hawaiite flow (47.9% SiO2, Table 2D; Figs. 4, 5A, and 6–8) is most like some early Miocene La Matancilla flows, because it has high La/Yb (32), Sm/Yb (7.6) and Ta/Hf (0.61) ratios.
EARLY MIOCENE MAGMATISM, INCIPIENT SHALLOW SUBDUCION UNDER THE NEUQUÉN BASIN, AND THE TECTONICS OF THE SOUTH-CENTRAL ANDES In the following discussion, the distribution and chemistry of the early to middle Miocene backarc volcanic rocks in the Neuquén Basin are combined with their structural setting to define three generalized magmatic episodes at ca. 24–20 Ma, 20–18 Ma, and 16–15 Ma. The petrologic, geochemical, and isotopic characteristics of the volcanic rocks in each episode are used to address differences in mantle and crustal source regions. These differences are subsequently incorporated into a tectonic model calling for: (1) incipient shallowing of the subducting Nazca plate north of the Cortaderas lineament after 20–19 Ma, (2) changes in interactions between South America and the plates in the Pacific Ocean to the west, and (3) changes in the absolute motion of South America relative to the mantle reference frame.
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Figure 6. (A) La/Yb ratio versus SiO2 (wt%), and (B) La/Sm versus Sm/Yb ratios for early Miocene Neuquén backarc volcanic rocks. La/Sm and Sm/Yb ratios, respectively, indicate the relative slopes of the light and heavy REE patterns in Figure 5. The general trends for decreasing melting percentages (lower f) of mantle spinel and garnet peridotites are shown. Points labeled PM and DMM are primitive mantle and depleted mid-ocean-ridge basalt (MORB) mantle values of Sun and McDonough (1989). Dates for Somuncura samples are from Kay et al. (1993, 2006). Other data sources and symbols are as in Figure 4.
Early Miocene Magmatic and Deformational Episodes in the Neuquén Basin The first magmatic episode produced the 24–20 Ma centers that predominantly erupted backarc alkali olivine basaltic lavas across the Neuquén Basin (Fig. 2A). The arrangement of these centers and their compositions are consistent with an association with extensional faults. The clearest faults have been identified in the Sierra de Chachahuén (Pérez and Condat, 1996; see Kay et al., this volume, chapter 10). Among the centers of this period are those that erupted the basalts exposed along the Filo Morado ridge in the Sierra de Huantraico and near Barda Costilla in the Sierra Negra (Fig. 3). The original morphology of these centers has been obscured by the subsequent contractional deformation that produced the Huantraico syncline. To the south, other monogenetic centers of this age erupted the Cerro Cabras basalts near the NWtrending Cortaderas lineament, while other to the east erupted the flows in the Sierra de Chachahuén and La Matancilla region. The second magmatic episode produced the ca. 20–18 Ma volcanic rocks, which outcrop in the southern and central Sierra de Huantraico. This activity appears to have been predominantly
Figure 7. (A) La/Ta ratio versus SiO2 (wt%) and (B) Ba/Ta versus La/Ta ratios for early Miocene Neuquén backarc volcanic rocks. High La/Ta, Ba/Ta, and Ba/La ratios are indicators of components associated with the subducting plate and overlying mantle wedge (arc values based on compilation of Hickey et al., 1986). Data for Cura Mallín volcanic rocks are from Burns (2002) and Kay et al. (this volume, ch. 2). Other data sources and symbols are as in Figure 4.
related to a large volcanic complex that produced the Huantraico sequence basaltic and hornblende-bearing mafic andesitic flows, agglomerates, sills, and dikes, the Pichi Tril andesites, and likely the Carrere tuffs (Fig. 3). The concentration of Pichi Tril andesites in the southeastern Sierra de Huantraico probably marks the principal eruptive center. This eruptive cycle is suggested to have started with explosive events that produced the Carrere tuffs, which were subsequently covered by basaltic to andesitic agglomerates and lava flows. The 18 ± 2 Ma Cerro Bayo de Huantraico can be interpreted as a dome and the ca. 19 Ma dikes and flows southeast of Cerro Bayo de Sierra Negra as coming from satellite centers. The proximity and chemical similarities of the dikes of the Desfiladero Negro unit of Ramos and Barbieri (1988) are consistent with their function as feeders to this volcanic complex and associated centers. The orientation of the dikes can be reconciled with the regional stress system (Cobbold and Rossello, 2003). In this framework, all of the units in the southern and central Sierra de Huantraico that were mapped with Eocene to late Miocene ages by Ramos and Barbieri (1988) and Rubinstein and Zappettini (1990) can be related to a ca. 20–18 Ma volcanic center. K-Ar ages are consistent with contemporaneous centers erupting north of the Sierra Negra, west of
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Figure 8. (A) Th/Hf versus Ta/Hf ratios for early Miocene Neuquén backarc volcanic rocks. Th/Hf ratios are measures of source components associated with a subducting slab and Ta/Hf ratios are measures of mantle enrichment. Data sources and symbols are as in Figures 4 and 7. (B) Chondrite-normalized La/Sm ratios (La = 0.378; Sm = 0.23) versus Th contents (ppm) for early Miocene Neuquén backarc basalts with >200 ppm Cr compared to calculated mantle melting models. Ticks on curves show percentages of partial melt. Curves show models for non-modal incremental batch melting of less enriched (50% midocean ridge basalt [MORB] plus 50% two times Leedy chondrite: La = 0.481, Sm = 0.380, Th = 0.060 ppm) and more enriched (La = 0.885, Sm = 0.414, Th = 0.113 ppm) sources (see Gorring and Kay, 2001). Relative melting percentages are minimums, since fractionation increases both Th concentrations and La/Sm ratios.
the Sierra de Chachahuén–La Matancilla region, and in the Cajón de Molle–Puntilla de Huincán–El Manzano region (Fig. 2) The third magmatic episode at ca. 16–15 Ma produced the andesitic Cerro Bayo de Sierra Negra center with its radial feeder dike system and the small 14.8 ± 1.5 Ma alkaline basaltic center at Cerro Villegas along the Cortaderas lineament (Fig. 3). Another flow that might be of this age is the 16 ± 5 Ma Cerro El Ramblón flow in the La Matancilla region (González Díaz, 1979). Existing ages are consistent with a gap in magmatism from ca. 19 Ma to 16 Ma and a subsequent break in backarc magmatism in northern Neuquén and southernmost Mendoza until the late Miocene (Kay et al., this volume, chapter 2). The rocks of all three magmatic episodes have been affected by contractional deformation. The question is the timing
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and duration of this deformation. A common view has been that of Kozlowski et al. (1996), who used regional stratigraphic correlations to argue that the principal contractional deformation ended in the Sierras de Huantraico and Negra between 16.3 ± 0.1 Ma and 6.7 ± 0.5 Ma. The older age limit comes from a K-Ar date on a deformed flow south of the Río Barrancas (southeast of Cerro Domuyo; Fig. 2A), and the younger limit, from a K-Ar date on an undeformed basalt flow overlying undeformed sediments (Tristeza Formation) south of Malargüe, Mendoza (Fig. 2). More recently, Ramos (1998) argued for compressional inversion of extensional structures in the Agrio fold-and-thrust belt contemporaneous with deposition of ~250 m of late Oligocene–early Miocene (?) (Tralahué) conglomerates (Fig. 2A). Cobbold and Rossello (2003) argued for deformation along the Cortaderas lineament contemporaneous with the formation of syndeformational growth strata in the Carrere Formation. If the Carrere Formation tuffs are early Miocene in age, these growth strata provide evidence for early Miocene deformation. Overall, contractional deformation in northern Neuquén and southern Mendoza seems to have begun by ca. 20 Ma as the style and distribution of volcanism changed across the region. Support for initial contractional inversion by 20–19 Ma also comes from seismic lines interpreted to show that a 1000-m-thick volcanic sequence that should be the early Miocene Huantraico sequence was affected by growth strata during the evolution of the Huantraico syncline (Viñes, 1990; see Cobbold and Rossello, 2003). Further support comes from amphibole phenocrysts in ca. 19 Ma Huantraico basaltic andesites. The instability of amphibole at low pressures requires that these magmas had a residence time in the middle to lower crust, as expected in a contractional regime that slows magma transit to the surface (e.g., Kay and Kay, 1994). Support for contraction after 20–19 Ma also come from elsewhere in the region (Fig. 1). To the west, Melnick et al. (this volume, chapter 4) argue that the main stage of Miocene normal faulting ended in the arc region by 18 Ma. To the south, Bechis (2004) presented evidence for compressional inversion of normal faults in the Nirihuau Basin before the eruption of the Collón Cura tuffs at 15 Ma. To the north, a gap in arc volcanism in the Coya Machalí Basin from 19 to 16 Ma is interpreted to coincide with compressional inversion of older normal faults (Kurtz et al., 1997; Godoy et al., 1999). Some like Radic et al. (2002) have used stratigraphic evidence to argue for extensional rifting until at least 13 ± 1.6 Ma in the southern Lonquimay Basin at ~38°S. A limited degree of local extension after 20 Ma is consistent with the alkaline chemistry of the 14.8 ± 1.5 Ma Cerro Villegas region basalt flow along the Cortaderas lineament. Source Regions of Early to Middle Miocene Neuquén Basin Backarc Magmas Chemical contrasts between 24 and 20 Ma flows and 20–15 Ma flows in the Neuquén Basin region are consistent with changes in both the composition of the mantle source and
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S.M. Kay and P. Copeland TABLE 3. Sr, Nd, AND Pb ISOTOPIC RATIOS FOR EARLY TO MIDDLE MIOCENE NEUQUÉN BACKARC VOLCANIC ROCKS 87 143 Sr/86Sr Nd/144Nd SiO2 Nd (wt%) (all are +) Sierra de Chachahuén and La Matancilla Regions DRC21 La Matancilla region – basalt flow at Puesto Juncal 47.8 0.7037114 0.5128531 4.2 DR19 Sierra de Chachahuén basalt flow 48.0 0.7040384 0.5128161 3.5 Sierra de Huantraico HDR18 Filo Morado olivine basalt flow HDR20 Huantraico olivine-clinopyroxene basalt flow HDR25 Huantraico – Desfiladero Negro basaltic dike HDR19 Huantraico hornblende-bearing andesite flow HDR14b Huantraico hornblende-bearing andesite flow
47.6 47.2 48.6 48.0 55.1
HDR20 HDR25 HDR14b
47.2 48.6 55.1
0.7038850 0.7032841 0.7032302 0.7033442 0.7036504 206
Huantraico olivine-clinopyroxene basalt flow Huantraico – Desfiladero Negro basaltic dike Huantraico hornblende-bearing andesite flow
Pb/204Pb 18.525 18.509 18.550
0.5128208 0.5128795 0.5128805 0.5128605 0.5128392 207
Pb/204Pb 15.580 15.561 15.583
3.6 4.7 4.7 4.3 3.9 208
Pb/204Pb 38.313 38.178 38.287
87 Sr/86Sr Sierra de Huantraico region† Huantraico sequence (Palaoco Formation) 0.70355 Huantraico sequence (Palaoco Formation) 0.70355 Pichi Tril Andesite 0.70361 Pichi Tril Andesite 0.70356 Parva Negra basalt 0.70351 Note: All data are modern-day ratios. Analyses were done at Cornell University. Analytical methods are discussed in Appendix 3. † Data from Ramos and Barbieri (1988).
the role of crustal contaminants as the tectonic regime became contractional. These changes are explored here by comparing the chemistry and isotopic compositions of the Neuquén backarc lavas to those of intraplate (OIB) and arc magmas (Figs. 6–10).
Figure 9. 87Sr/ 86Sr ratios versus εNd values for early Miocene Neuquén backarc volcanic rocks and other late Oligocene–Miocene volcanic rocks between 32°S and 42°S. Data for samples from: (1) the Coya Machalí arc are from Kay et al. (2005), (2) the Colbún region are from López-Escobar and Vergara (1997), (3) the Somuncura Province and the Ventana arc are from Kay et al. (1993, 2006), and (4) arc and forearc volcanic rocks between 37° and 42°S are from Muñoz et al. (2000). Data for Neuquén backarc volcanic rocks are from Table 3. Note that samples with ages from ca. 24–20 Ma are more enriched in 87Sr/ 86Sr and more depleted in εNd than samples with ages from 19 to 16 Ma. See text for further discussion.
An Intraplate-Like Mantle Source for 24–20 Ma Magmas The chemical characteristics of the 24–20 Ma volcanic rocks are typical of intraplate (OIB) mafic magmas that have mantle sources nearly devoid of subduction-type components (Figs. 5A). This is shown by: (1) enriched high field strength element (HFSE) concentrations relative to the light REEs, as indicated by low La/Ta (<14) ratios (Figs. 7A and 7B), (2) a lack of evidence for subducted crustal material, as shown by low Th/La ratios (0.09–0.13; Table 2B–2C; <0.12 in midocean-ridge basalt [MORB]–OIB lavas—see Klein and Karsten, 1995), and (3) a lack of evidence for excess fluidmobile hydrous elements, as indicated by low Ba/La ratios (<16; Fig. 7B). In detail, the Chachahuén–La Matancilla basalts, which are those farthest from the trench, show the least evidence of subducted components (La/Ta < 10; Ba/La < 15, and Th/La < 0.11) . Marginally higher La/Ta (up to 14), Ba/La (up to 19), and Th/La (up to 0.14) ratios (Figs. 6–8) in Filo Morado basalts are consistent their more proximal position to the trench. The paucity of evidence for slab components in all of these magmas is consistent with their having similar 87Sr/ 86Sr ratios that plot at lower ε Nd values when compared with Oligocene to early Miocene Ventana, Coya Machalí, and Colbún arc volcanic rocks farther west (Fig. 9). Simple trace-element models for 24–20 Ma basalts with >250 ppm Cr that are based on the steepness of the heavy REE patterns (Fig. 6B) as well as La/Sm ratios and Th concentra-
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Figure 10. Plots of 207Pb/ 204Pb and 208Pb/ 204Pb versus 206Pb/ 204Pb for early Miocene Neuquén Basin backarc volcanic rocks and other late Oligocene–Miocene samples between 32°S and 42°S (see Fig. 1). Data sources are as in Figure 9. Data from Chon Aike rhyolite are from Gorring and Kay (2001). MORB—mid-ocean-ridge basalt; SVZ—Southern Volcanic Zone; NHRL—Northern Hemisphere reference line.
tions (Fig. 8B) are consistent with these magmas having been derived from ~3%–5% partial melting of a garnet-bearing peridotite source. Higher La/Sm ratios at a given Th concentration and higher Ta/Hf ratios (0.46–0.67 vs. 0.38–0.53) in Chachahuén–La Matancilla basalts than in Filo Morado basalts (see Figs. 8A and 8B) are expected with a more enriched intraplate-like mantle source farther from the trench. Entry of a Subduction-Related Component into the Mantle Source after ca. 20 Ma Evidence for a change in the backarc mantle source comes from chemical characteristics in younger than 20 Ma volcanic rocks that begin to approach those in subduction zone magmatic rocks. The more arc-like chemistry of these volcanic rocks is signaled by: (1) a relative depletion of HFSE, as shown by lower TiO2 contents (1.2% to 1.7%) and higher La/Ta (15–21) ratios (Fig. 7), and (2) excess fluid-mobile elements, as shown by generally higher Ba/La (most 15–22; Fig. 7B) and lower Nb/U ratios (Table 2). Evidence for a hydrous component also comes
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from amphibole phenocrysts that appear in basaltic andesitic and andesitic lavas. Other indicators of a change in the mantle source are lower La/Sm ratios at a given Th content (Fig. 8B), and lower Ta/Hf ratios (0.24–0.42; Fig. 8A) in the 20–15 Ma magmas. The even stronger arc-like tendencies in the Molle sequence volcanic rocks are shown by their generally lower TiO2 contents (1.0–1.5%) and higher La/Ta (17–20), Ba/La (19–25), and Th/La (0.11–0.25) ratios. In accordance with a stronger subduction zone influence on a more depleted mantle source, the Molle sequence samples have the lowest Ta/Hf ratios (0.21–0.27; Fig. 8A) among the 24–15 Ma lavas. Several indicators point to higher partial mantle melting percentages for the 20–15 Ma magmas than for the older than 20 Ma magmas, which is consistent with hydrous fluxing of the mantle source leading to higher backarc magma production rates after 20 Ma. The first indicator is lower Th concentrations at a given La/Sm ratio ((Fig. 8B), which fits with derivation of the Huantraico magmas by up to ~7% partial melting and the Molle magmas by up to 4% partial melting of a depleted mantle source. The second indicator is flatter heavy REE patterns in the ca. 20–15 Ma basalts (La/Yb = 5–11 and Sm/Yb = ~2–2.8; Figs. 5 and 6) that point to either higher melting percentages leaving less residual garnet in the mantle source or shallower melting depths within the spinel peridotite stability field than for the magmas older than 20 Ma. Evidence for a change in the mantle magma source near 20 Ma also comes from higher 87Sr/ 86Sr ratios and lower εNd values in samples with ages greater than 20 Ma than in those younger than 20 Ma (Fig. 9; Table 3). These isotopic differences are hard to explain by a common contaminant within the crust, since basalts with ages greater than 20 Ma have higher 87Sr/ 86Sr ratios (>0.7037) and lower ε Nd values (<+4.2) than an andesite (55% SiO2) younger than 20 Ma. Relating these isotopic differences solely to regional changes in the basement that might be inferred from the distribution of the samples is inconsistent with consistent temporal trends in trace element ratios across the region. Higher K/Cs ratios in samples younger than 20 Ma (K/Cs = >40000 versus <28000, Table 2) than in samples older than 20 Ma point to a higher proportion of crustal contaminants in the younger magmas. A Role for Crustal Contaminants in Younger than 20 Ma Magmas Chemical differences between magmas younger and older than 20 Ma indicate incorporation of more or different crustal components into the magma source at ca. 20 Ma. These crustal components can be introduced in two ways: (1) into the mantle magma source by sediment subduction and/or forearc subduction erosion processes, and (2) by contamination of mantlederived magmas within the crust. Evidence for the first process comes from Th/La ratios that are closer to the MORB-OIB range (0.9–0.12) in basaltic samples older than 20 Ma and closer to crustal-like values (as high as 0.17) in basaltic samples younger than 20 Ma (see Table 2). Evidence for entry of a crustal compo-
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nent as the younger basaltic magmas evolved to trachyandesiticdacitic magmas in the crust comes from a general trend to more upper crustal–like incompatible trace-element and isotopic ratios in the more silicic samples (see following). In general, the characteristics of the Pichi Tril trachyandesites support a genetic relationship with the Huantraico basalts and andesites, as shown by overlapping fractionationindependent incompatible trace-element ratios like La/Ta, Ba/Ta, Th/Hf, and Ta/Hf (see Figs. 5–8). Higher La/Yb ratios in the more evolved magmas, which are largely attributable to steeper light REE slopes and lower middle REEs, fit with fractionation of clinopyroxene and amphibole. Relatively small negative Eu anomalies (Eu/Eu* > 0.80) and high Sr (>520 ppm) concentrations in the more Si-rich magmas (Table 2; Figs. 5B and 5C) show that plagioclase fractionation was not a dominant factor. These conditions fit with fractionation of hydrous magmas near the base of the crust where clinopyroxene and amphibole crystallize before plagioclase (e.g., Moore and Carmichael, 1998). The presence of large clinopyroxene phenocrysts in the basaltic lavas also fits with fractionation of hydrous basaltic magmas in the deep crust (e.g., Sisson and Grove, 1993). The best evidence for within-crustal components increasing with SiO2 content in the younger than 20 Ma magmas comes from systematic changes in isotopic, Th/La, and Nb*/U (where Nb = Ta × 17; see Sun and McDonough, 1989) ratios. Isotopic support comes from higher 87Sr/ 86Sr and 207Pb/ 204Pb ratios and lower εNd values in the andesite than the basalts as shown in Figure 9 (e.g., 87Sr/ 86Sr = ~0.7033 vs. ~0.7036) and Figure 10. As such, one contaminant seems to be a Sr-rich cumulate or crustal-melt residue that is compatible with high Sr contents in trachyandesitic-dacitic magmas (>520 ppm, up to 900 ppm). A likely contaminant is the deep crustal residue of the widespread Permian-Triassic Choiyoi granite-rhyolite complex (Kay et al., 1989). Support for an additional upper-crustal contaminant comes from the progressive decrease in Nb*/U ratios, which are as low as 28 in the basalts to as low as 12 in the trachyandesites. For comparison, Nb*/U ratios range from ~38 to 58 in OIB-MORB lavas to <20 in continental crust. A similar argument can be made from Th/La ratios that range from 0.11 to 0.17 in the mafic lavas to 0.15–0.25 in the trachyandesites (Tables 2C and 2D) and that compare with ratios ranging from <~0.12 in OIB-MORB lavas to >0.22 in continental crust (OIB-MORB and continental crustal Nb*/U and Th/La ratios are as in Klein and Karsten, 1995). A Model for Early Miocene Shallowing of the Subduction Zone beneath the Neuquén Basin A straightforward way to explain a change to a contractional stress regime in conjunction with introduction of a subduction component into the backarc mantle magma source in the early Miocene is to call upon shallowing of the subducting oceanic plate beneath the Neuquén Basin at 20–19 Ma. Alternative explanations like the eastward advance of a steeply sub-
ducting oceanic slab can be ruled out by the appearance of subduction-related components in younger than 20 magmas that erupted more than 200 km east of the arc magmatic front. A lithospheric-scale cartoon cross section for the early Miocene evolution of the Neuquén Basin north of the Cortaderas lineament is shown in Figure 11. A cross section at ca. 22 Ma (Fig. 11A) shows the situation as arc-like basaltic to dacitic magmas (Figs. 7 and 8; Burns, 2002; Kay et al., this volume, chapter 2) were erupting in the Cura Mallín intra-arc basin, and backarc olivine alkaline basalts were erupting in the Sierras de Huantraico and Negra and the La Matancilla–Sierra de Chachahuén region. A period of mild extension in the arc has been proposed to explain the features of the Cura Mallín intra-arc basin (Suárez and Emparan, 1995; Burns, 2002; Jordan et al., 2001). The associated Cura Mallín volcanic rocks have the chemistry expected of low- to medium-K arc magmas (Burns, 2002; Kay et al., this volume, chapter 2) erupting at or near the arc front. At the same time, magmas erupted in the backarc show an intraplate (OIB-like) chemistry, with virtually no evidence for a subducted component in their mantle source. Their composition, volume, association with monogenetic or simple polygenetic cones, and distribution are in accord with passage though a thin crust in a mildly extensional regime in which mantle-generated melts easily transited to the surface. As a whole, the arc chemical signature of the Cura Mallín volcanics, the near absence of subduction-type components in the Filo Morado basalts, and the lack of any subductiontype component in the Sierra de Chachahuén–La Matancilla region basalts require a steep subduction zone at this time. The cross section at ca. 19 Ma (Fig. 11B) shows a shallower subduction zone to account for the introduction of subducted components into the backarc magma source and the termination of the extensional stress regime across the region. Evidence for the end of extension in the arc region comes from the shutdown of the Cura Mallín intra-arc basin at this latitude (Burns, 2002; Burns et al., this volume, chapter 8). Evidence in the backarc comes from: (1) termination of eruption of olivine alkali basalts from widespread monogenetic centers, (2) eruption of a range of basaltic to hornblende-bearing trachyandesitic-dacitic magmas with weak subduction signatures from large volcanic complexes in the Sierras de Huantraico and Negra and in southern Mendoza, and (3) the formation of the Huantraico syncline and associated contractional structures. Support for a hydrated backarc mantle magma source and deep-crustal backarc magma evolution prior to eruption in a contractional regime comes from evidence for crystallization of clinopyroxene and hornblende at the expense of olivine and plagioclase, and for entry of crustal contaminants into evolving magmas. Contractional deformation in the Sierra de Huantraico and adjoining backarc could have been facilitated by hydraulic weakening of the overlying continental lithosphere by slab-derived fluids (e.g., James and Sacks, 1999). Subduction erosion of forearc crust and lithosphere as the subducting plate shallowed could have facilitated the introduction of crustal lithosphere into the backarc mantle wedge.
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Figure 11. Lithospheric scale cartoon cross sections across the Neuquén Basin near 37°S for the early to middle Miocene illustrating the model for changing slab dip discussed in the text. Arrows indicate a change from extensional to contractional stress at ca. 20 Ma. Magmatic centers with symbols in gray have arc or backarc geochemical signatures, whereas those in black have intraplate geochemical signatures. Percentage melt is an average estimate of degree of mantle partial melting. See text for discussion.
Larger-Scale Plate Tectonic Processes and the Evolution of the Neuquén Basin Two related fundamental questions are why early Miocene mafic alkaline magmatism initiated across the backarc in the Neuquén Basin contemporaneous with the formation of the Cura Mallín intra-arc basin (e.g., Vergara et al., 1997; Jordan et al., 2001) and why the stress regime across the region became contractional around 20 Ma. Understanding the events in the Neuquén Basin requires taking a broader view of late Oligocene–early Miocene magmatism and forearc, intra-arc, and backarc basin formation in the south-central Andes (region between 33°S and 45°S in Fig. 1). The relative motion of South American–Pacific oceanic plate interactions and the motion of the South American plate over the underlying mantle also need to be considered. South American–Oceanic Plate Interactions Most authors relate the tectonic setting of the Oligocene to early Miocene magmas and basins in this region to upper- and lower-plate adjustments associated with the breakup of the Farallon plate during Chron 6 (Handschumacher, 1976) near ca. 23.5 Ma (Cande and Kent, 1992). This breakup occurs in conjunction with a large jump in the rate (~7–15 cm/yr) and an oblique to a nearly normal directional change in the convergence of South America relative to the prebreakup Farallon and the postbreakup Nazca plates (e.g., Pardo Casas and Molnar, 1987;
Somoza, 1998). A series of pull-apart models have been invoked to explain the formation and distribution of the early Miocene Nirihuau, Cura Mallín, and Coya Machalí basins (e.g., Cazau et al., 1987; Suárez and Emparan, 1995; Godoy et al., 1999). Although the distribution of these basins fits an oblique pull-apart stress regime, Burns and Jordan (1999) and Jordan et al. (2001) rejected this model citing a lack of evidence for concurrent strikeslip faulting and an apparent contradiction with a nearly normal Nazca–South America convergence regime. They argued that intra-arc basin formation could be explained by continental lithospheric doming and collapse as arc magmatism reinitiated above a rapidly subducting oceanic plate in the aftermath of plate reorganization. Jordan et al. (2001) argued that the change to a compressional regime at 20–18 Ma could be explained by the convergence regime reaching a stable condition at that time. In contrast, Muñoz et al. (2000) and Burns (2002) called for slab roll-back models to explain extension in forearc basins south of 36°S and the Cura Mallín intra-arc basin, respectively. A factor in understanding Andean margin dynamics during this time could be in the details of the interactions among the oceanic plates to the west. These changes can be essentially invisible in South American convergence models based on a few snapshots of magnetic anomalies. To explore these relationships, the tectonic model for the 28.3–16 Ma evolution of the southeastern Pacific Ocean in Figure 12 from Tebbens et al. (1997) and Tebbens and Cande (1997) is compared with events along the Andean margin in Figure 12. Important features in the
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Figure 12. Model for the tectonic evolution of southeast Pacific oceanic crust between ca. 28 and 16 Ma from Tebbens and Cande (1997) compared to magmatic events between 42°S and 32°S on the South American plate and postulated motion of South America relative to the underlying mantle. On oceanic model, note rotation of oceanic spreading centers between 25.8 and 20 Ma, formation of microplates (MP) that transfer parts of one oceanic plate to the other, and northward migration of Pacific–Farallon-Nazca–Antarctica triple junction from 33 to 16 Ma. Light lines are troughs (Tr.); black lines are fracture zones; heavy black lines are fossil ridges; dark regions are transferred microplates; dashed line shows track of triple junction. See Figure 1 for localities on the South American plate.
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Slab shallowing and the westward drift of South America oceanic model include northward migration of the Pacific– Antarctic–Farallon/Nazca triple junction, changes in orientations of spreading ridges, and the formation of microplates during ridge jumps. Factors of potential importance to the Andean margin are the change from a margin-oblique trend for the NWtrending Pacific-Farallon ridge at 25.8 Ma to the nearly marginparallel trends for the Nazca-Pacific and Nazca-Antarctic ridges at 20 Ma. These changes require: (1) clockwise rotations of 5° for the Pacific-Antarctic ridge, 38° for the Nazca-Antarctic ridge, and 10° for the Pacific-Nazca, and (2) compensating ridge jumps to create the Selkirk and two other possible microplates. Importantly, the period of ridge rotations and the associated creation of microplates overlaps the formation of the early Miocene Andean extensional forearc, intra-arc, and backarc basins. The major effect of the oceanic plate interactions would be small progressive changes in relative South American– oceanic plate-convergence parameters. Such rotations could facilitate basin formation that would cease at ca. 20 Ma as the Pacific-Nazca and Nazca-Antarctic ridges became nearly parallel to the Chile Trench. Motion of South America Relative to the Mantle Hotspot Reference Frame Another factor to consider is the motion of the South American plate relative to the underlying mantle in the hotspot reference frame. Support for Oligocene to early Miocene changes in the absolute motion of South America relative to the underlying mantle comes from studies of both the South American and African plates. In particular, Silver et al. (1998) argued that the most recent Andean deformation cycle that began at ca. 25 Ma was driven by an increase in the velocity of the South American plate over the underlying mantle as the African plate slowed down. Since then, O’Connor et al. (1999) have demonstrated that the relative rate of motion of Africa compared to hotspots has been 20 ± 1 mm/yr since 19 Ma, as opposed to 30 ± 3 mm/yr before 45 Ma. There are no good constraints on the rate from 45 to 19 Ma. As such, a major change in the westward rate of South America relative to the hotspot reference frame at ca. 19 Ma is a possible contributor to a change from an extensional to a contractional regime along the Andean margin. Other changes in the rate of South America relative to the underlying mantle could be signaled by an upsurge of hotspot volcanism on the African plate at ca. 25 Ma; an increase in the initiation of African plate oceanic hotspots between 19 and 30 Ma, and the ca. 29–26 Ma Somuncura eruption. Roll of Absolute Mantle Motion in the Late Oligocene to Early Miocene Evolution of the Neuquén Basin and Surrounding Region The possible role of the absolute motion of South America relative to the hotspot reference frame in the evolution of the Neuquén Basin and the south-central Andes is examined here in light of the regional magmatic and deformational history and
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isotopic signatures of the erupted magmas (Figs. 9 and 10). The data are shown to fit with a model for a nearly fixed position for South America relative to hotspots in the late Oligocene, a relative eastward retreat of South America in the early Miocene, and a relative westward advance of South America in the late early Miocene. Late Oligocene The late Oligocene evolution of the Neuquén region is marked by a magmatic quiescence that is contemporaneous with the eruption of alkaline pre-plateau magmas followed by voluminous ca. 29–26 Ma tholeiitic plateau magmas in the Somuncura region at 39° to 43°S (Fig. 1; e.g., Ardolino and Franchi, 1993, Kay et al., 1993, 2004). Kay et al. (1993) used a lack of evidence for backarc extension and chemical analogies with oceanic-island basalt lavas to argue that the Somuncura plateau magmas formed in response to instabilities in the mantle convection system just before the breakup of the Farallon plate. Subsequently, Muñoz et al. (2000) related the plateau lavas and Oligocene to Miocene volcanic rocks to the west with ages generally older than 28 Ma and younger than 24 Ma (Fig. 1) to a slab-window created by slab roll-back during the breakup of the Farallon plate. Further data to consider are the high 87Sr/ 86Sr ratios and low εNd values of the plateau lavas compared to other Oligocene to Miocene volcanic rocks in the region (Fig. 9). In analogy with models for mafic volcanic rocks formed over a stationary mantle in Africa (see summary by O’Connor et al., 1999), the relatively enriched isotopic ratios and large volumes of plateau lavas and a contemporaneous lack of backarc extension can be reconciled with the plateau magmas forming in a hot asthenospheric mantle at a time when South America was relatively stationary over the hotspot reference frame. Early Miocene until 19 Ma The subsequent development of the Neuquén region up until ca. 19 Ma occurred in an extensional regime associated with the eruption of the backarc alkali olivine basalts in the Neuquén Basin, the Somuncura post-plateau lavas, and backarc basalts elsewhere in the region (e.g., Máquinas basalts near 32°S; Kay et al., 1991), as well as the formation of backarc, intra-arc, and forearc basins (Nirihuau, Lonquimay, Cura Mallín, Coya Machalí, and marginal basins in Fig. 1). An episode of slab roll-back (retreat) as called upon by Muñoz et al. (2000) for the Valdivia and nearby forearc basins or Burns (2002) for the Cura Mallín intra-arc basin seems to best explain the tectonic picture at this time. In detail, slab roll-back at a time of eastward advance of South America over the underlying mantle provides the only clear explanation as to why an extensional regime extended more than 500 km from the trench at a time of relatively steep and rapid subduction of the Nazca plate. The nature of basin development and magmatism fits with combining changing Nazca plate-convergence directions in Figure 12 with eastward retreat of the South American plate. A broader
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expanse of magmatism and basin development in the south (Fig. 2) fits with a weaker continental lithosphere in this region in the aftermath of the Somuncura magmatic event. This tectonic picture is consistent with the isotopic data in Figure 9. Lower 87Sr/ 86Sr and higher εNd values in Somuncura post-plateau than plateau lavas can be attributed to a diminished influence of a deep asthenospheric source and an increased influence of a depleted arc-like mantle source. The lower 87Sr/ 86Sr ratios and higher εNd values in ca. 23–20 Ma Neuquén Basin backarc lavas can be explained by even less influence from a deep asthenospheric source. The still higher εNd values at a given 87Sr/ 86Sr ratio in Ventana, Coya Machalí, and Colbún arc magmas fit an even more depleted arc-type mantle source modified by Sr-bearing fluids from the subducting slab. The large range of values in the 37° to 42°S arc-forearc magmas can be explained by complex mixing of arc and deep-mantle sources. Early to Mid-Miocene after 19 Ma Early Miocene extension ended at ca. 20 Ma, coincident with cessation of magmatism in the forearc basins, structural inversion of forearc, arc, and backarc basins, and introduction of subduction components into the backarc magma source north of ~37.5°S. As suggested by Jordan et al. (2001), the time of change to compression might simply reflect the stabilization of South American–Nazca plate interactions after the breakup of the Farallon plate. More recently, Yañez and Cembrano (2005) have argued that the Miocene extensional regime was coincident with a period of weak coupling between the Nazca and South America plates and that a compressional regime emerged as coupling increased in response to both a decreasing convergence rate and subducted plate age along the margin. A problem with this model in the Neuquén region is that their calculations show that the degree of plate coupling has changed little in the last 25 m.y. in this region (their Fig. 7). Other factors related to the cessation of extension could be that the offshore oceanic ridges reached a parallel alignment with the Chile Trench (Fig. 12) and that the initiation of compression coincided with one of the three major periods of Tertiary plate reconfiguration (at Chron C6 [o] in Fig. 12) in the southeast Pacific (Tebbens and Cande, 1997). Another factor could be the initiation of trench advance at a time of change to near trenchnormal westward-directed motion and an increase in the velocity of the westward advance of South America over the underlying mantle. A decrease in 87Sr/ 86Sr ratios and an increase in εNd values in older than 20 Ma Neuquén region backarc magmas fit with a lessening of the influence of an Atlantic-like asthenospheric mantle as it was progressively cut off as the South American plate overrode the subarc mantle and a shallower subducting plate (Fig. 11). Apparent relative magmatic gaps or lulls from ca. 18 to 16 Ma that correspond to an apparent ca. 19 to ca. 16 Ma gap in Neuquén Basin backarc volcanism can be seen in existing radiometric age dates for magmatic rocks across the south-central Andean region (Cazau et al., 1987; López Escobar and Vergara, 1997;
Muñoz et al., 2000; Kay et al., 2005). A compressional phase from ca. 20–19 to ca. 16 Ma fits with rapid trench advance during a rapid period of relative westward motion of South America over the mantle reference frame. CONCLUSIONS Relatively abrupt changes at ca. 24 and ca. 20 Ma in backarc magmatic and structural styles along the south-central Andean margin are seen in the Neuquén Basin between 36° and 37.6°S. The change at ca. 24 Ma coincides with the sudden eruption of mafic OIB-type alkali olivine basalts in a mildly extensional setting, and that at ca. 20 Ma, with the eruption of more arc-like basaltic to dacitic magmas in a contractional setting. The time span of alkaline magmatism overlaps the formation of extensional-setting–related ca. 24–20 Ma sedimentaryvolcanic forearc, intra-arc, and backarc basins all along the margin, and the change near 20 Ma coincides with contractional inversion of these basins, cessation of forearc magmatism, and magmatic changes that require arc-type components entering the backarc mantle north of ~37.5°S. This entire period coincides with the rapid, nearly normal Nazca–South America plate convergence regime that emerged after the breakup of the Farallon plate. The ca. 24–20 Ma extensional period occurred during a time when oceanic-ridge jumps formed oceanic microplates as the oceanic ridges between the Nazca, Pacific, and Antarctic plates rotated into near parallelism with the Chile Trench. The onset of the compressional period coincided with a somewhat lower South America– Nazca convergence rate. Arguably, the most important control on changes in magmatic and structural styles in the Neuquén Basin and the southcentral Andes during this time was the relative velocity and direction of motion of the South American plate over the hotspot reference frame. The extensional tectonic regime and the mantle sources of the erupted magmas from 24 to 20 Ma fit with slab roll-back related to the relative eastward motion of South America over the mantle. The discontinuous nature of the early Miocene basins and backarc mafic eruptions fit with trench-oblique motion of South America with respect to the underlying mantle. In the same way, the post–20 Ma change to a contractional regime fits with accelerated westward motion of South America over the underlying mantle. The introduction of arc components into the backarc mantle in the Neuquén Basin, north of ~37.5°S latitude, concurrent with an incipient shallowing of the subducting slab is compatible with this model. A change to a contractional regime near 20 Ma fits with a model linking the slow-down of the African plate with a speed-up in the western motion of South America over the mantle (Silver et al., 1998) and evidence for a decreased rate of motion for Africa in the last 19 m.y. (O’Connor et al., 1999). Confirmation of trench roll-back preceding trench advance awaits better constraints on the relative motion of South America over the mantle. Also needed are better structural constraints on the nature, mechanisms, and extent of early Miocene extension.
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Slab shallowing and the westward drift of South America ACKNOWLEDGMENTS Funding for this study was provided by a grant from Repsol YPF petroleum company with additional support from U.S. National Science Foundation (NSF) grant 00-87515. The authors thank REPSOL-YPF for funding and permission to publish the results from: Kay, S.M., Tertiary to Recent Magmatism and Tectonics of the Neuquén Basin between 36.5°S and 38°S latitude, Final Report to REPSOL-YPF, April 30, 2001, 215 pp. Oscar Mancilla (REPSOL-YPF) is thanked for support during the course of the study, Linda Godfrey for assistance with isotopic analyses, Daniel Ragona for participating in the field work and helping with trace-element analyses, Matthew Burns for participating in the field work, and Facundo Fuentes for processing the Thematic Mapper images. Discussions with R.W. Kay, T. Zapata, D. Ragona, M. Burns, and V.A. Ramos were valuable in interpreting the results. The manuscript was improved by helpful comments from Adam Goss, an informal review by Jason Phipps Morgan, and formal reviews by Peter Cobbold and Constantino Mpodozis. APPENDIX 1. SAMPLE DESCRIPTIONS AND LOCALITIES Sierra Negra Region HDR1. Basaltic andesite along road just southwest of Cerro Bayo de Sierra Negra, 37°23.0′S, 69°24.8′W. HDR2. Clinopyroxene-bearing basaltic dike near road southwest of Cerro Bayo de Sierra Negra, near 37°23′S, 69°23′W. HDR3. (a) Fine-grained andesite and (b) Coarser grained dacite; along seismic line just off road near Cerro Bayo de Sierra Negra, 37°23.2′S, 69°21.74′W.
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HDR14. Hornblende-bearing basaltic andesite block in large slump from main cliff above puesto west of Puesto Gonzalez: (a) lower columnar jointed flow, (b) upper, more massive flow, 37°35.8′S, 69°42.13′W. HDR15. Basaltic andesite flow just west of yellow house at puesto, vesicle-rich, 37°32.20′S, 69°42.4.7′W. HDR16. Basaltic andesite flow northwest of puesto near Cerro Sarmiento, 37°28.76′S, 68°40.88′W, 1650 ± 50 m. HDR17. Las Cabras Basalt, small outcrops north of Cerro Villegas, 37°31.32′S, 69°52.22′W. HDR18. Olivine basalt east of Filo Morado near well platform, 37°15.76′S, 69°39.13′W, 917 ± 25 m. HDR19. Clinopyroxene-bearing basalt flow along San Jorge company road, east side of Huantraico anticline, 37°23.98′S, 69°30.93′W, 1239 ± 30m. HDR20. Clinopyroxene-bearing basalt flow along San Jorge company road, east side of Huantraico anticline, reddish outcrop, 37°24.96′S, 69°31.97′W, 1172 ± 34 m. HDR21. Clinopyroxene-bearing basalt flow, dirt track just north of puesto at Cerro Los Cerrillos, Chapúa basalt on map, 37°24.31′S, 69°37.70′W, 1222 ± 27m. HDR22. Basaltic agglomerate, past small lake on eastward extension of seismic line east of east-west straight segment of Huantraico road, 37°30.80′S, 69°36.72′W, 1286 ± 27 m. HDR23. Basaltic dike with large clinopyroxene phenocrysts just south of Route 7, west of Parva Negra, 37°38.18′S, 69°35.83′W, 873 ± 33 m. HDR24. Basaltic dike with large clinopyroxene phenocrysts, Cerro Loma Viejo dike, 37°31.33′S, 69°26.82′W, 842 ± 35 m. HDR25. Basaltic dike with large clinopyroxene phenocrysts, Desfiladero Negro dike, 37°35.09′S, 69°28.57′W, 781 ± 40 m. Sierra de Chachahuén
Sierra de Huantraico Region TDR14. Olivine basalt flow at road junction from Buta Ranquil to Rincón de Las Sauces, 37°10.50′S, 69°38.70′W. TDR32. Olivine basalt flow at end of dirt track in quebrada, south of road to Rincón de Las Sauces, north end of Sierra Negra, near Barda de Castillo, 37°12.61′S, 69°29.22′W, 882 ± 39 m elevation. HDR4. Olivine basalt near Puesto Cerro Pampa de Las Liebres, 37°33.8′S, 69°50′W. HDR6. (a) Basaltic clast in agglomerate, (b) dike cutting agglomerate, (c) basaltic dike, (d) andesitic clast in tuffs; base of Bajada El Toro; 37°31.24′S, 69°38.02′W, 1283 ± 22 m. HDR7–HDR10. All in Cerro Bayo de Huantraico region near 37°31′S, 69°33′W, 1286 ± 27 m. HDR7. Basaltic flow near magnetite-apatite deposit. HDR8. Basalt associated with tuffs above HDR7. HDR9. Basaltic dike cutting agglomerate section. HDR10. Andesite of Cerro Bayo de Huantraico. HDR11. Basaltic dike on road to puesto (outpost), 37°31.6′S, 69°32′W. HDR13. Basaltic dike near Parva Negra flow, 37°41.8′S, 69°34′W.
DR19. Olivine basalt flow along road in quebrada near Puesto Cortaderita, Fe-stained bands in outcrop, 37°3.72′S, 68°51.74′W, 1396 ± 40 m. DR58. Olivine basalt flow south of Puesto Zuñiga, 37°05′44.4?S, 68°54′36.6?W. DRC4. Olivine basalt flow from Arroyo Los Golpeados area, edge of quebrada near outcrop of Cretaceous Agrio Formation, 37° 05.44′S, 68° 54.72′W. Hoja La Matancilla—North of Sierra de Chachahuén DRC17. Basalt flow just east of Cerro Las Lajas, along seismic line southwest of quarry, 36°50.2′S, 69°04.07′W. DRC18. Basalt flow along side of road near El Ramblón, 3 km east of Las Lajas, 36°49.62′S, 69°1.03′W, 1226 m. DRC20. Basalt next to Puesto Peligroso, steep outcrop just north of Puesto: (a) Fresh coarse-grained columnar flow, upper basalt, (b) altered with carbonate, lower basalt, 36°31′S, 68°32′W. DRC21. Basalt flow, southwest of road to Puesto Junial, near Cerro El Azufre, 36°36.8′S, 68°35.5′W.
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APPENDIX 2. 40Ar/ 39Ar AGE SPECTRUM
Figure A1. Age spectra are shown for samples with ages in Table 1 dated by the 40Ar/ 39Ar step-heating method (Mitchell, 1968) at the University of Houston in the laboratory of Peter Copeland. Analytical details can be found in Jordan et al. (2001).
APPENDIX 3: ANALYTICAL METHODS Major-element chemical analyses were performed on fused glasses using the JEOL 733 Superprobe in the Cornell University Materials Science Center (CCMR). Samples with <60% SiO2 were melted in a molybdenum strip furnace in an Ar atmosphere, whereas samples with >60% SiO2 were mixed with a meta-borate flux and melted in carbon crucibles in air at 1000 °C for 30 min in a muffle furnace. Microprobe analyses were preformed at 15 kV
with a current of 15 A using wavelength-dispersive spectrometers. Reported analyses are averages of five 20–30-μm-diameter spots. Smithsonian standard Juan de Fuca glass was used as a secondary standard. Analyses on fluxed glasses were normalized to 100%. Trace-element analyses were done by instrumental neutron activation analyses (INAA) at the Ward Reactor Center at Cornell University. Powdered samples (~0.5 g) were sealed in high-purity silica glass tubes and irradiated in the Cornell Triga reactor at a neutron flux of 5 × 1013 neutrons cm–2 s–1 for 2 h. Gamma-ray
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Slab shallowing and the westward drift of South America spectra were acquired using an ORTEC intrinsic Ge coaxial detector (20% efficiency; resolution of 0.7 KeV at 1.33 MeV), a Nuclear Data ADC, and a multichannel analyzer. Samples were counted for a minimum of 2 h (up to 10 h), 7 and 40 d after irradiation. Data were reduced using the Cornell data reduction program. Corrections for peak interferences were made on Ce (for Fe), Nd (for Br), Tb (for Th), Eu (for Ba), Lu (for U), and Yb (for Th). Whole-rock FeO concentrations were used as internal flux monitors and trace-element concentrations are proportional to FeO concentrations. FeO concentrations based on INAA counting data and Na2O concentrations determined by INAA data were crosschecked against whole-rock analyses to ensure analyses were accurately matched. Priority was given to Na2O concentrations from INAA analyses. Reported analyses are volatile free, because they were proportioned to Fe concentrations from volatile-free analyses on glasses. See Kay et al. (1987) for further discussion. All isotopic analyses were done at Cornell University on a VG Sector thermal ionization mass spectrometer (TIMS). Sr isotopic ratios were measured on W single filaments using a quadruplecollector dynamic procedure. Ratios were normalized to an 87Sr/ 86Sr ratio of 0.1194. Nd isotopic analyses were measured on single Re filaments using a quintuple-collector dynamic procedure. Average analytical values for 87Sr/ 86Sr for NBS987 were 0.710221 (±0.000044) and 143Nd/ 144Nd ratios for La Jolla were 0.511888 (±0.000055). 143Nd/ 144Nd ratios were corrected to a value of 0.5118624 for the La Jolla standard, and εNd was calculated based on a value of –15.15 for La Jolla. Pb isotopic ratios were corrected for mass fractionation based on ratios of 206Pb/ 204Pb = 16.931, 207Pb/ 204Pb = 15.485, and 208Pb/ 204Pb = 36.681 measured on Pb standard NBS SRM981. REFERENCES CITED Ardolino, A., and Franchi, M., 1993, El vulcanismo Cenozoico de la Meseta de Somuncura, Río Negro y Chubut, in XII Congreso Geológico Argentino (Buenos Aires): Actas, v. 4, p. 225–235. Ardolino, A., Franchi, M., and Fauqué, L., 1996, Geología y recursos minerales del Departamento Añelo, Province Neuquén map: Buenos Aires, Programa Nacional de Cartas Geológicas de la República Argentina, Subsecretaria de Minería de la Nación, Servicio Geológico, Añales 25, escala 1:250,000. Baker, P.E., Rea, N.J., Skarmeta, J., Caminos, R., and Rex, D.C., 1981, Igneous history of the Andean Cordillera and Patagonian Plateau around latitude 46°S: Philosophical Transactions of the Royal Society London, v. A303, p. 105–149. Baldauf, P.E., 1997, Timing of the uplift of the Cordillera Principal, Mendoza Province, Argentina. [PhD. thesis]: Washington, D.C., George Washington University, 356 p. Bechis, F., 2004, Estudio geológico y estructural de la región media de los Ríos Nirihuau y Pichileufu, Provincia del Río Negro [Trabajo final de Licenciatura]: Buenos Aires, Argentina, Universidad de Buenos Aires, 120 p. Burns, W.M., 2002, Tectonic and depositional evolution of the Tertiary Cura Mallín basin in the southern Andes (36.5 to 38°S lat.) [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, Goettingen, Germany: Paris, Insitut de Recherche pour le Developpement, p. 115–118.
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MANUSCRIPT ACCEPTED BY THE SOCIETY 22 DECEMBER 2005
Printed in the USA
Geological Society of America Special Paper 407 2006
Evolution of the late Miocene Chachahuén volcanic complex at 37°S over a transient shallow subduction zone under the Neuquén Andes Suzanne Mahlburg Kay* Institute for the Study of the Continents and Department of Earth and Atmospheric Sciences, Snee Hall, Cornell University, Ithaca, New York 14853, USA Oscar Mancilla* Repsol YPZ, Buenos Aires, Argentina Peter Copeland* Department of Geosciences, University of Houston, Houston, Texas 77204, USA
ABSTRACT The evolving chemistry of the Chachahuén volcanic complex provides evidence for transient entry of a subduction zone component into the mantle wedge over a late Miocene shallow subduction zone under the Neuquén Basin. The Chachahuén complex, which is in the backarc of the Andean Southern Volcanic Zone near 37°S and some 500 km east of the Chile Trench, occurs at the intersection of NE and SE fault systems that parallel regional trends. Support for a shallow subduction-zone setting at the time of eruption and during the contractional uplift of the Sierra de Chachahuén comes from K/Ar and new 40Ar/ 39Ar ages, mineral assemblages, major and trace element chemistry, and Nd-Sr-Pb isotopic compositions. Importantly, the chemistry of the Chachahuén rocks requires an arc-like component in the mantle that is absent in both early Miocene or Pliocene alkaline lavas erupted in the same region. The oldest Chachahuén volcanic rocks are the ca. 7.3–6.8 Ma Vizcachas group orthopyroxenebearing andesites to rhyodacites that erupted from fissures and small centers along the NE-trending fault system. Intraplate chemical tendencies in the most silicic samples are attributed to mantle-derived basalts interacting with a lower crust that has a chemical imprint that reflects older alkaline magmatic events. Younger Chachahuén group volcanic rocks erupted at ca. 6.8–6.4 Ma from vents generally aligned along the NE-trending fault system and ca. 6.3–4.9 Ma magmas that erupted from a trap-door– type caldera and flanking stratovolcanoes along the NW-trending fault system. These high-K basaltic to dacitic rocks contain amphibole phenocrysts and show arc-like high field strength element depletions that are the strongest in basaltic andesite lavas. Parallels between Chachahuén volcanic rocks and uplift of the Sierra de Chachahuén with late Miocene Pocho volcanic rocks and uplift of the Pampean Ranges over the modern Chilean flat-slab support transient Miocene shallow subduction zone under the Neuquén Basin. *E-mails:
[email protected];
[email protected];
[email protected].
Kay, S.M., Mancilla, O., and Copeland, P., 2006, Evolution of the late Miocene Chachahuén volcanic complex at 37°S over a transient shallow subduction zone under the Neuquén 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. 215–246, doi: 10.1130/2006.2407(10). For permission to copy, contact
[email protected]. ©2006 Geological Society of America. All rights reserved.
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S.M. Kay et al. Keywords: Andean arc magmatism, shallow subduction, Neuquén Basin, Miocene, caldera complex, trace elements, isotopes.
INTRODUCTION Neogene magmatism has been common in the far backarc of the Andean Southern Volcanic Zone, north of the NWtrending Cortaderas lineament at ~38° to 37°S. Among the volcanic centers that erupted more than 450 km east of the Chile Trench in the Neuquén Basin (Figs. 1 and 2) are the latest Miocene to early Pliocene Chachahuén volcanic complex near 37°S and the Cerro Nevado-Plateado volcanic complexes near 35.5°S (Bermúdez, 1991). The petrologic characteristics of these dominantly hornblende-bearing basaltic andesitic to dacitic complexes stand in sharp contrast to those of early Miocene alkaline olivine basaltic flows (Kay and Copeland, this volume, chapter 9) and post-Miocene alkaline flows in the same region. The post-Miocene flows are those associated with the extensive Llancanelo, Payún Matrú, and Auca Mahuída alkaline volcanic fields (Bermúdez and Delpino, 1989; Stern et al., 1990; Saal et al., 1993; Saal, 1994; Kay et al., this volume, chapter 2) that cover a vast region of the backarc (Fig. 2). Muñoz et al. (1989) and Bermúdez et al. (1993) noted the contrast between the eruption of the Chachahuén, Cerro Nevado and Plateado complexes in uplifted crustal blocks and the younger alkaline flows predominantly in lower-lying regions. The arc-like petrologic features of the Chachahuén (Kay, 2001; Kay and Mancilla, 2001) and Cerro Nevado-Plateado (Bermúdez, 1991) volcanic complexes in the Neuquén Basin are particularly notable for their resemblance to those of late Miocene to Pliocene volcanic rocks that erupted in the blockfaulted Pampean Ranges, up to 700 km east of the Chile Trench over the Chilean flat-slab to the north (Fig. 1). Importantly, the sources of the Pampean magmas have been associated with shallowing of the subducting Nazca plate (Kay and Gordillo, 1994; Kay and Mpodozis, 2002). In contrast, the latest Miocene to early Pliocene Neuquén Basin volcanic centers occur east of a now steeply dipping segment of the Benioff zone (e.g., Cahill and Isacks, 1992). This raises the question as to whether their arc-like features can be related to a shallowly dipping subducting slab whose geometry has since changed. Initial support for a change in slab dip beneath the Neuquén Basin comes from a compositional study of Pliocene to Holocene primitive basaltic volcanic rocks erupted between 35°S to 37°S at distances of 50–250 km east of the arc front. In this study, Saal et al. (1993) and Saal (1994) argued that the following temporal trends were best explained by a small increase in the subduction rate and/or steepening of the subduction angle of the Nazca plate: (1) hypersthene to nepheline normative compositions, (2) lower to higher incompatible trace element contents, (3) more to less arc-like incompatible trace element
ratios, (4) higher to lower Sr and Pb isotopic ratios, and (5) lower to higher εNd. They correlated these and other trends with a decrease in slab fluid flux, an increase in depth of melt segregation, and a decrease in mantle partial melting with increasing distance from the trench. In this context, the purpose of this paper is to characterize the volcanological, petrological, and geochemical characteristics of the Chachahuén volcanic complex and to interpret these features in a model for transient Miocene to Pliocene shallow subduction of the Nazca plate beneath the Neuquén Basin. In doing this, the evolution of the ca. 7.3–4.8 Ma Chachahuén complex is evaluated in light of new field work and reinterpreted map patterns (Table 1; Fig. 3A–C), the four new 40Ar/ 39Ar and 24 previously unpublished K-Ar ages from Pérez and Condat (1996) in Table 2, 65 new whole-rock major- and trace-element analyses (Table 3), and new Sr, Nd, and Pb isotopic analyses on seven samples (Table 4). These data are used to argue that the Chachahuén volcanic complex evolved from (1) small vents erupting silicic magmas with intraplate to arc-like chemistry along faults, to (2) a nested caldera and “trap-door” type caldera complex erupting basaltic to hornblendebearing andesitic to dacitic magmas with arc-like chemistry. Mantle-derived magmas are argued to have interacted with a crust that initially had an intraplate-like chemistry, but was progressively modified by the intrusion of arc-like mantlederived magmas. The volcanic rocks with the most arc-like chemical features are shown to be mafic flows from small latestage peripheral vents. The magnitude of the shallowing of the Nazca plate is evaluated by making comparisons with the modern along-strike geometry of the Nazca plate as well as models for the evolving geometry of the Nazca plate over the Chilean flat-slab region near 30°S. The case for transient shallowing of the Nazca plate under the Neuquén Basin is supported by chemical contrasts with early Miocene and post–middle Pliocene intraplate-like alkaline basalts. REGIONAL SETTING OF THE CHACHAHUÉN VOLCANIC COMPLEX The Chachahuén volcanic complex in the Sierra de Chachahuén is the only late Miocene volcanic center in the central part of the Neuquén Basin (Figs. 1 and 2). The late Miocene volcanic rocks are best exposed in the central high parts of the range. Outcrops at lower elevations are in the midst of late Pliocene to Pleistocene basaltic lava cones and associated flows (Fig. 2). The first reconnaissance studies of the Chachahuén volcanic complex were made in conjunction with the 1:200,000 regional-scale geologic mapping of the Sierra de Chachahuén (Holmberg, 1962) and La Matancilla (González Díaz, 1979)
Evolution of the late Miocene Chachahuén volcanic complex
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Figure 1. Topographic relief map (Shuttle Radar Topographic Mission image; http://www2.jpl.nasa.gov/srtm/) of the south central Andean arc and backarc showing the location of the Chachahuén volcanic complex in the Sierra de Chachahuén relative to the Chile Trench, the Holocene Southern Volcanic Zone arc, and the late Miocene-Pliocene Cerro Nevado-Plateado, San Luis, and Pocho volcanic fields. The Pocho and San Luis fields are in the block-faulted Pampean Ranges over the volcanically inactive Chilean flat-slab region, whose southern limit is approximately indicated by the dashed line. The box shows the general area of Figure 2. The lines labeled A and B show the locations of the cross sections in Figure 11.
sheets. The volcanic rocks were assigned Eocene to Miocene ages based on field relations and late Miocene K-Ar ages, which are compiled in Linares and González (1990). Subsequently, when the Sierra de Chachahuén became a target for oil exploration, Pérez and Condat (1996) produced a 1:10,000scale map of the range for the YPF petroleum company. They proposed the volcanic stratigraphy shown in Table 1 by correlating unconformities between measured sections, and reported the 24 late Miocene K-Ar ages listed in Table 2. More recently,
Kay (2001) and Kay and Mancilla (2001) emphasized the petrologic features of the Chachahuén volcanic center that make it unique among the Neogene volcanic centers in this part of the Neuquén Basin. On a regional scale, the location of the Chachahuén volcanic complex, like other Neogene magmatic centers and structural features in the Neuquén Basin, can be tied to older extensional structures that controlled the Mesozoic evolution of the basin (e.g., Vergani et al., 1995; Zapata et al., 1999; Zencich,
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Figure 2. Geologic map of Tertiary to Holocene magmatic rocks in southern Mendoza and northern Neuquén Provinces showing the Chachahuén and Plateado-Nevado volcanic complexes, the distribution of early Miocene and post–early Pliocene alkaline volcanic rocks, and the Cortaderas lineament, which marks the southern limit of Miocene to Holocene backarc volcanism. Map is based on the 1:750,000-scale geologic maps of Mendoza (1993) and Neuquén (1995) Provinces published by the Secretaría de Minería, Dirección Nacional del Servicio Geológico, Buenos Aires, Argentina.
2000). In particular, the Chachahuén complex appears to be at the intersection of long-lived NW-SE– and NE-SW–trending fault systems that were most recently reactivated by late Miocene compressional and Pliocene to Holocene extensional deformation (Fig. 3B). Within the Sierra de Chachahuén, a major NW-trending fault system extending through the Rio Café to the southeast is consistent with morphological and volcanic features (see Figs. 3 and 4). Segments of this system were mapped by Pérez and Condat (1996). An intersecting NE-
trending system parallel to the Arroyo los Golpeados is consistent with the distribution of Cretaceous sedimentary and early Miocene basalt outcrops in the Sierra de Chachahuén. As seen on provincial geologic maps of Mendoza and Neuquén, similar NW-SE–trending faults and fractures are widespread in the region. One of the most important NW-trending features is the Cortaderas lineament (Ramos, 1978), which generally defines the southern limit of Neogene backarc magmatism in the Neuquén Basin (Fig. 2).
Evolution of the late Miocene Chachahuén volcanic complex
Age
Group (this paper)
TABLE 1. STRATIGRAPHY AND VOLCANIC UNITS OF THE SIERRA DE CHACHAHUÉN Perez and Condat (1996) Thickness Description (m) Basaltic cones and lava flows
Quaternary–latest Pliocene Earliest Pliocene Late Chachahuén Cerro Chachahuén Flows latest Miocene (youngest part) until ca. 4.9 Late Miocene Late Chachahuén Unit 7 ca. 6.3 to 5.4 Ma (includes part of unit 5, all of units 6 and 7) Unit 6
Late Miocene ca. 6.9–6.3 Ma
219
Early Chachahuén (includes most of unit 3, unit 4, and most of unit 5)
Late Miocene ca. 7.3–7.0 Ma
Vizcachas
Early Miocene ca. 24–20 Ma
Chachahuén Matancilla flows
Cretaceous
Neuquén Group
~90
Four- to six-meter-thick packets of thin, dark to reddish-gray aphanitic mafic lava flows interlayered with medium gray pyroclastic flows. Peripheral lava flows.
~80
Small areal distribution.Top: Irregular gray, brown columnar jointed mafic andesite lava flows intercalated with pyroclastic flows. Fine conglomerates and sandstones at base. Top: dark gray, columnar jointed andesitic flows associated with course interbedded pyroclastic flows. Abrupt changes in facies and thickness. Base: Thick sequence of light gray to white tuffs, tuffaceous "mass flows – avalanche deposits, debris flows" interbedded with volcanic agglomerates. Thick basal sequence of greenish-gray mafic andesitic lavas with thick interlayered pyroclastic flow and fall deposits, minor conglomerates, wedges of sandstone. Thick sequence dominated by light-colored fall & pyroclastic flows. Also volcanic conglomerates, mafic andesite lava flows, and epiclastic deposits. Variable thickness pumiceous pyroclastic flows at base. Gray volcanic conglomerates interlayed with remnant small pyroclastic flows. Sandy conglomerates. Intervals of pyroclastic flows interbedded with minor lava flows. Surge deposits with irregular and chaotic geometry. Large blocks in conglomerate. Basaltic flows interbedded with medium-grained sandstones containing basaltic clasts. Conglomerates, marine fossils, chert and laminated quartz sandstones.
>150
Unit 5
~170
Unit 4
140
Unit 3
>120
Unit 2, part of unit 3
1–100
Angular unconformity Unit 1
0–70
Angular unconformity Red and white sandy conglomerates, red and green tuffaceous shales
Volcanic Divisions, Age, and Chemistry of the Chachahuén Volcanic Complex The distribution, style, and chemical signatures of the principal late Miocene volcanic units in the Sierra de Chachahuén are summarized in this section. The stratigraphic units of Pérez and Condat (1996) are combined into the Vizcachas, Early Chachahuén and Late Chachahuén groups, as shown in Table 1. These groups are based on reinterpreting the Pérez and Condat (1996) units using concepts of volcanic stratigraphy, satellite imagery (Fig. 4), the K-Ar ages of Pérez and Condat (1996), the new 40Ar/ 39Ar ages in Table 2, and the chemical and isotopic analyses in Tables 3 and 4 and plots in Figures 5–10. Descriptions and localities of the analyzed samples are in Appendix 1; spectra associated with 40Ar/ 39Ar dating are in Appendix 2; analytical methods are described in Appendix 3. Maps of the Vizcachas, Early Chachahuén, and Late Chachahuén groups in Figure 3 show the Pérez and Condat (1996) units incorporated in these groups. In general, the Vizcachas group includes most of unit 2 (Fig. 3A) and parts of unit 3 (not shown on Fig. 3A), the Early Chachahuén group (Fig. 3B) includes the majority of unit 3, all of unit 4, and part of unit 5, and the Late Chachahuén group (Fig. 3C) includes part of unit 5, all of units 6 and 7, the Cerro Chachahuén flows, and scattered peripheral flows. These divisions reconcile the inconsistencies that Pérez and Condat (1996) encountered in correlating their map units and K-Ar ages (see Tables 1 and 2). All ages are listed in Table 2 and plotted either on the maps in Figure 3 or on the TM image in Figure 4. The new groups are discussed individually in the following sections.
Late Miocene Vizcachas Group The Vizcachas group consists of the oldest late Miocene volcanic rocks in the Chachahuén volcanic complex. Included in this group are the Eocene(?) volcanic rocks of Holmberg (1962) in the Cerro de Las Vizcachas region and the late Miocene volcanic rocks in the Sierra de Chachahuén mapped as unit 2 and part of unit 3 by Pérez and Condat (1996). The Cerro de Las Vizcachas region units include dacitic dikes cutting porphyritic diorites and Mesozoic sedimentary rocks along with andesitic to dacitic ignimbrites and lava flows. All are considered to be late Miocene in age based on a new 40Ar/ 39Ar age of 7.80 ± 0.60 Ma from the groundmass of a Cerro de Las Vizcachas region rhyodacitic ignimbrite DRC24a (Table 2). The majority of the Vizcachas group outcrops in the Sierra de Chachahuén are concentrated in a line from the Puesto Zuñiga region to northeast of Cerro Bayo (Fig. 3A). These rocks are typically variable-thickness silicic pyroclastic flows, volcaniclastic sandstone and surge deposits, block and ash deposits, and very subordinate lava flows. A particularly distinctive feature is the presence of silicic vitrophyres concentrated in the Cerro Bayo region. A late Miocene age for the Vizcachas group in the Sierra de Chachahuén is confirmed by 40Ar/ 39Ar ages of 7.90 ± 0.3 Ma for plagioclase and 7.28 ± 0.07 Ma for the groundmass of a clast (DRC7c in Table 2) in an ignimbrite east of Puesto Zuñiga. The groundmass age is within the uncertainty of a K-Ar age of 7.3 ± 0.4 Ma in Pérez and Condat (1996) for a unit 2 andesite along a fault in the Arroyo de los Golpeados (Table 2). Other K-Ar ages in Table 2 from Pérez and Condat (1996) that limit the age of the Vizcachas group are 7.1 ± 0.4 Ma for a tuff in Cerro Corrales, 7.0 ± 0.4 Ma for a mafic lava flow
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Figure 3. Maps showing distribution of the volcanic rocks of (A) late Miocene Vizcachas group, (B) late Miocene Early Chachahuén group, and (C) late Miocene–earliest Pliocene Late Chachahuén group relative to proposed eruptive centers (dashed ovals) and faults (solid lines) and K-Ar and 40Ar/ 39Ar dates (Table 2). Map units are based on those in Table 1 from Pérez and Condat (1996) with the modifications discussed in the text. Eruptive centers and fault systems (dashed lines) are shown on Thematic Mapper (TM) images in Figure 4 for comparison.
north of Cerro Cabeza, and 7.0 ± 0.4 Ma for a mafic intrusive in the Cretaceous Agrio Formation near Cerro Corrales. Chemical and isotopic analyses of Vizcachas group samples are listed in Table 3A and plotted in Figures 5–10. Cerro de Las Vizcachas region samples come from welded ignimbrites (DR44 and DRC24A) and a dacitic porphyry (DRC23). Sierra de Chachahuén samples come from a glassy margin of a dike (DR16) and vitrophyres in unit 2 and 3 near Cerro Bayo (DR20, DR2, DRC8, and DR55), from block and ash deposits near Cerro Bayo (DR15) and the head of Arroyo los Golpeados (DRC3), and from pumices and clasts in unit 2 pyroclastic
flows near Puesto Zuñiga and at the base of Cerro Boina (DRC7a to c and DR23). Most of the samples are orthopyroxene-bearing dacites to rhyodacites (63%–70% SiO2) that plot near the mid-K to high-K field border for arc rocks (Fig. 5A), have Na2O plus K2O concentrations in the trachyandesite range, and fall in the calc-alkaline field on a FeO/MgO versus SiO2 diagram (Fig. 5B). Their distinctive trace-element features can be seen in Figures 6A and 6C. They include very high Ba (900–1250 ppm) and Sr (650–1100 ppm) contents, steep rare earth element (REE) patterns (La/Yb = 20–42) characterized by very steep light (La/Sm ~ 6–11) and moderate heavy (Sm/Yb ~ 3–4; Fig. 7) REE slopes,
Evolution of the late Miocene Chachahuén volcanic complex
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TABLE 2. 40Ar/39Ar AND K/Ar AGES FOR CHACHAHUÉN VOLCANIC COMPLEX Type
Sample
Latitude °S
Vizcachas Group 40 Ar/39Ar DRC24a 37°05 06.0 South of Cerro Corrales 40 Ar/39Ar DRC7c 37°04 20.4 K/Ar K/Ar K/Ar K/Ar
Ch-8 Ch-3 Ch-54 Ch-11
37°04 37°04 37°00 37°04
Early Chachahuén Group Cerro Bayo Region K/Ar Ch-58 37°02 K/Ar Ch-50 37°02 South of Cerro Corrales K/Ar Ch-7 37°05 K/Ar Ch-22b 37°05 K/Ar Ch-20a 37°07 K/Ar Ch-5 37°05 K/Ar Ch-24b 37°05 Cerro Cabeza Region K/Ar Ch-57 37°00 Cerro Bayo Region K/Ar Ch-30 37°04 K/Ar Ch-39 37°03 K/Ar Ch-52 37°04 Cerro Cabeza Region K/Ar Ch-57 37°00
Longitude °W
Group, Center, and Description of Locality
Age (Ma)
Type
69°00 06.0
Cerro de Las Vizcachas – welded rhyodacitic tuff.
7.80 ± 0.60
Groundmass
68°53 52.2
Dacitic clast in ignimbrite east of Puesto Zuñiga
46.2 3.2 20.8 40.6
68°55 68°55 68°53 68°54
Head of A° Los Golpeados along fault in unit 2, “intrusive.” Base of “unit 6” in la Buitrera, Cerro Corrales, andesitic tuff. North of Cerro Cabeza – unit 2 basaltic andesite lava flow. Basaltic andesite intrusive in Cretaceous Agrio Formation (DR29).
7.90 ± 0.3 7.28 ± 0.07 7.3 ± 0.4 7.1 ± 0.4 7.0 ± 0.4 7.0 ± 0.4
Plagioclase Groundmass Andesite Whole rock Whole rock Whole rock
25.5 36.4
68°50 45.4 68°51 35.7
Top of Cerro Ratón – unit 3? Andesite flow? North side of Cerro Bayo – andesite “intrusive.”
6.9 ± 0.3 6.7 ± 0.3
Whole rock Whole rock
40.9 22.3 07.3 30.4 37.2
68°55 68°57 68°57 68°54 68°54
Andesite intrusive into fault in Cretaceous Neuquén Group. Hill east of Puesto Abandonado, andesite dike. Basaltic andesite next to Cerro Panul. Right side Arroyo de Los Golpeados andesite “intrusive” near DRC6. Andesite on right side of Arroyo de Los Golpeados near DRC3.
6.8 ± 0.3 6.5 ± 0.3 6.5 ± 0.3 6.4 ± 0.3 6.1 ± 1.3
Whole rock Whole rock Whole rock Whole rock Whole rock
24.7
68°52 22.5
NE flank of Cerro Cabeza – unit 3 andesite flow.
6.5 ± 0.4
Whole rock
49.6 47.0 02.7
68°52 39.3 68°51 17.1 68°52 56.1
Northwest of Puesto J.A. Sosa, basaltic andesite “intrusive.” Andesite flow on north side of Cerro Chachahuén “unit 3.” Salida Agua del Torduco, “unit 3” south of Cerro Boina, andesite “intrusive.”
6.6 ± 0.3 6.4 ± 0.3 6.3 ± 0.3
Whole rock Whole rock Whole rock
24.7
68°52 22.5
NE flank of Cerro Cabeza - unit 3 andesite flow.
6.5 ± 0.4
Whole rock
11.7 19.2 01.6 20.3
41.9 07.0 48.0 53.8 40.3
Late Chachahuén Group K/Ar Ch-26b 6.4 ± 0.3 Whole rock 37°00 35.3 68°56 44.0 Pyroxene andesite flow, base of Cerro El Tanque in Río de Café, unit 6. K/Ar Ch-1 6.3 ± 0.3 Whole rock 37°04 06.1 68°54 47.1 Base of “Unit” 6, columnar jointed andesite flow, Cerro Corrales (like DR26). K/Ar Ch-2 6.1 ± 0.3 Whole rock 37°04 05.2 68°55 19.2 Mafic andesite dike in La Buitrera, west of Cerro Corrales (like DR27). K/Ar Ch-6 5.9 ± 0.3 Whole rock 37°05 31.3 68°55 21.1 Mafic andesite flow, base of “Unit 5”, west of Arroyo de Los Golpeados. K/Ar Ch-19a 5.9 ± 0.3 Whole rock 37°05 57.9 68°53 34.3 West of Cañadón del Camino – andesite debris flow? (like DR49). K/Ar Ch-33 6.3 ± 0.3 Whole rock 37°04 28.3 68°50 55.5 Mafic andesite flow at top of Cerro Chachahuén like DR10. 6.1 ± 0.3 Whole rock K/Ar Ch-40b 37°01 22.0 68°54 22.4 Andesite in middle of Quebrada Fiera, west of Cerro Cabeza, unit 3. 6.0 ± 0.3 Whole rock K/Ar Ch-59 37°05 28.2 68°48 20.8 South of Cerro Bombilla, andesite “intrusive.” K/Ar Ch-32 North side of Cerro Chachahuén, andesite “intrusive” like DR12. 5.6 ± 0.3 Whole rock 37°04 22.3 68°5 22.0 40 Ar/39Ar DRC13 4.85 ± 0.03 Groundmass 68°51 14.4 Flow from Cerro Chachahuén, southeast of Cañadon del Camino. 37°06 00.0 Note: Spectra for 40Ar/39Ar ages are in Appendix 2. Ages on samples with label beginning with Ch are K/Ar whole-rock ages from Perez and Condat (1996). DR and DRC designations in parentheses are for samples in Table 3 from the same or nearby location as the dated sample.
small negative Eu anomalies (Fig. 6C), and La/Ta and Ta/Hf ratios (Figs. 8 and 9) that range from intraplate-like in Cerro de Las Vizcachas region rhyodacites (La/Ta ~ 11 and Ta/Hf ~ 1) to more arc-like in Sierra de Chachahuén dacites (La/Ta ~ 13–23 and Ta/Hf ~ 0.4–0.7). High arc-like Ba/La (28–46) and Ba/Ta ratios (Fig. 8) reflect high Ba contents. The highest 87Sr/ 86Sr (0.7040–0.7045) and lowest 206Pb/ 204Pb (18.55) ratios along with the lowest εNd (+0.8 to +2) in the Chachahuén volcanic complex are in the Vizcachas group (Table 4; Fig. 10). The most extreme values are in a rhyodacite with 70.7% SiO2. Early Chachahuén Group The Early Chachahuén group includes the rest of the volcanic rocks in units 2–4, part of unit 5, and some of the “intrusives” of Pérez and Condat (1996). Most of this group consists of unit 3 and 4 volcanic rocks that are widely distributed in a
NE-SW–trending belt across the Sierra de Chachahuén (Fig. 3B). Those in unit 3 largely consist of variable-thickness pyroclastic deposits that include ignimbrites, lavas flows, volcanic conglomerates, and sandstones. Those in unit 4 consist of a thick sequence of pyroclastic flows and fall deposits associated with sparse lava flows and conglomerates. The unit 5 rocks in this group overlie unit 4 in the southwesternmost area of the map. Most of unit 5, which is dominated by greenish-gray mafic andesitic lava flows intercalated with pyroclastic fall and flow deposits and conglomerates, is included in the younger Late Chachahuén group. The “intrusives,” like all of those mapped by Pérez and Condat, (1996), have glassy or devitrified groundmass consistent with their forming as extrusive or shallow intrusive domes, porphyries, or welded ignimbrites. K-Ar ages reported by Pérez and Condat (1996) for the Early Chachahuén group as constituted here range from 6.9 ± 0.3 to 6.3 ± 0.3 Ma (Fig. 3B; Table 2).
30.0 52.7 18.9 3.56 0.93 0.392 1.05 0.150 979 1050 6.9 4.0 9.9 3.0 1.6 9.0 15 11 12
32.2 54.4 19.4 3.72 0.99 0.415 1.13 0.162 980 1023 7.8 3.9 9.9 3.3 1.7 10.6 13 10 14
31.8 54.0 17.5 3.08 0.82 0.272 0.88 0.112 800 1012 8.0 4.4 10.7 3.1 1.8 6.8 16 11 9
31.6 56.5 12.9 2.95 0.71 0.243 0.87 0.143 798 1029 8.4 4.5 11.3 3.1 2.0 5.8 15 5 8
Na2O + K2O 7.77 7.74 7.15 8.22 6.96 7.21 7.49 7.32 FeO/MgO 3.25 2.42 1.82 2.08 1.95 2.04 1.90 2.12 La/Sm 10.2 10.9 7.0 7.9 8.4 8.7 10.3 10.7 Sm/Yb 4.2 3.0 3.3 3.0 3.4 3.3 3.5 3.4 La/Yb 42.2 33.0 23.3 23.5 28.5 28.6 36.1 36.2 Eu/Eu* 1.01 1.06 0.90 0.94 0.92 0.93 1.01 0.93 Ba/La 46.2 44.3 38.1 38.5 35.0 31.7 31.9 32.6 Ba/Ta 480 488 522 542 640 595 548 508 La/Ta 10.4 11.0 13.7 14.1 18.3 18.7 17.2 15.6 Th/La 0.006 0.37 0.30 0.35 0.33 0.31 0.34 0.36 Ta/Hf 1.06 0.92 0.72 0.63 0.55 0.53 0.60 0.65 Th/Hf 4.33 3.79 2.94 3.15 3.33 3.03 3.48 3.62 K/Cs 3072 3036 4807 2867 3148 2515 2942 2923 Th/U 1.8 1.9 1.7 2.1 2.5 2.5 2.5 2.5 Latitude °S 37°5.76 37°05.10 37°05.10 37°2.94 37°3.72 37°2.94 36°59.52 37°02.36 Longitude °W 69°0.24' 68°52.01' 68°51.74' 68°52.01 68°53.29' 68°52.11 69°00.1 69°00.1 Elevation (m) 1650 ± 38 1396 ± 40 1650 ± 38 1450 ± 30 1607 Note: All sample locations and descriptions are in Appendix 1. Analytical methods are discussed in Appendix 3.
31.5 54.7 18.5 3.99 1.05 0.430 1.34 0.213 830 1210 8.4 5.2 11.1 3.5 2.2 7.0 11 10 11 7.45 1.95 11.1 3.5 38.6 1.06 32.7 511 15.7 0.35 0.70 3.89 2596 2.4 37°2.04 68°50.78' ~1530
33.0 55.7 14.9 2.98 0.78 0.217 0.86 0.136 818 1078 8.4 5.0 11.7 3.0 2.1 5.4 16 8 7 7.46 3.42 4.3 4.6 19.9 0.87 38.0 894 23.5 0.29 0.30 2.06 5711 2.6 37°05.44 68°54,72 ~1320
30.1 57.5 25.4 7.00 1.17 0.526 1.51 0.217 1037 1144 3.8 3.3 8.8 4.3 1.3 7.4 11 8 10
7.97 3.18 8.5 3.4 28.4 0.98 37.8 565 15.0 0.34 0.63 3.22 3292 2.3 37°04.34 68°53.87 1464 ± 34
32.8 57.4 19.4 3.88 1.00 0.331 1.16 0.150 846 1240 6.7 4.9 11.2 3.5 2.2 6.2 33 20 9
6.97 1.87 10.6 3.3 35.1 1.09 27.6 499 18.1 0.35 0.47 2.94 2647 4.3 37°04.34 68°53.87 1464 ± 34
33.0 66.7 15.5 3.12 0.85 0.233 0.94 0.113 1097 910 7.4 2.7 11.5 3.9 1.8 4.2 20 15 7
7.59 2.03 9.2 3.5 32.1 0.84 34.6 714 20.6 0.31 0.42 2.68 3277 2.6 37°04.34 68°53.87 1464 ± 34
32.6 58.6 23.2 3.53 0.79 0.319 1.01 0.158 995 1129 7.5 3.9 10.0 3.7 1.6 4.1 17 11 6
La Ce Nd Sm Eu Tb Yb Lu Sr Ba Cs U Th Hf Ta Sc Cr Ni Co
26.1 47.6 16.8 3.73 0.93 0.373 1.12 0.161 658 993 4.4 4.6 7.8 2.6 1.9 8.3 28 18 12
26.9 44.2 14.4 2.65 0.65 0.180 0.64 0.092 671 1245 7.4 5.8 10.5 2.4 2.59 2.9 13 10 3
SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O P2O5 Total 26.2 45.6 13.8 2.40 0.62 0.168 0.79 0.113 638 1161 7.6 5.0 9.8 2.6 2.4 3.3 14 11 4
TABLE 3A. WHOLE-ROCK CHEMISTRY: CHACHAHUÉN VOLCANIC COMPLEX, VIZCACHAS GROUP Cerro de Las Vizcachas Sierra de Chachahuén Cerro Bayo Region Arroyo East of Puesto Zuñiga west south west north northeast northeast Golpeados middle base top DR44 DRC24a DRC23 DR16 DR20 DR15 DR2 DRC8 DR55 DRC3 DRC7a DRC7b DRC7c ignimbrite porphyry dike vitrophyre clast vitrophyre vitrophyre vitrophyre ignimbrite clast pumice clast 70.73 70.70 63.86 62.59 63.76 63.53 66.14 67.77 68.41 62.87 65.09 66.19 66.82 0.16 0.20 0.63 0.52 0.66 0.70 0.47 0.38 0.36 0.69 0.46 0.34 0.38 15.85 15.86 16.64 16.59 16.88 16.36 16.58 16.51 15.94 17.80 16.56 17.43 17.13 1.65 1.84 4.29 4.13 4.16 4.61 3.25 2.77 2.68 4.62 3.65 2.85 2.65 0.15 0.05 0.07 0.11 0.08 0.09 0.08 0.08 0.08 0.08 0.08 0.10 0.06 0.51 0.76 2.36 1.99 2.13 2.26 1.71 1.31 1.38 1.35 1.15 1.53 1.31 2.86 2.91 4.91 5.31 5.17 5.12 4.11 3.94 3.66 4.85 4.78 4.48 3.87 5.01 4.97 4.61 5.33 4.34 4.85 4.64 4.37 4.81 4.85 5.32 4.63 4.61 2.76 2.77 2.54 2.89 2.62 2.36 2.85 2.95 2.64 2.61 2.65 2.35 2.98 0.02 0.05 0.17 0.21 0.13 0.15 0.05 0.09 0.07 0.21 0.19 0.09 0.13 99.69 100.11 100.07 99.65 99.93 100.03 99.88 100.16 100.03 99.93 99.92 99.98 100.03
8.11 2.20 9.9 3.4 34.0 1.05 31.7 646 20.4 0.31 0.42 2.69 3202 3.4 37°03.3 68°52.3 ~1580
38.5 72.9 22.5 3.90 1.11 0.378 1.13 0.181 1279 1221 7.1 3.5 12.0 4.5 1.9 4.6 17 14 8
DR23 pumice 65.07 0.41 17.02 3.09 0.17 1.41 4.36 5.35 2.75 0.19 99.82
Cerro Boina
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TABLE 3B. WHOLE-ROCK CHEMISTRY: CHACHAHUÉN VOLCANIC COMPLEX EARLY CHACHAHUÉN GROUP (SILICIC)
SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O P2O5 Total La Ce Nd Sm Eu Tb Yb Lu Sr Ba Cs U Th Hf Ta Sc Cr Ni Co Na2O + K2O FeO/MgO La/Sm Sm/Yb La/Yb Eu/Eu* Ba/La Ba/Ta La/Ta Th/La Ta/Hf Th/Hf K/Cs Th/U Latitude °S Longitude °W Elevation (m)
Cerro de la Gloria DR35 DR34 intrusive intrusive 66.08 68.82 0.44 0.33 16.79 16.74 3.11 2.80 0.16 0.08 0.89 0.45 4.03 2.34 4.73 4.34 3.66 3.97 0.14 0.06 100.03 99.94 35.6 71.4 29.4 5.53 1.33 0.666 2.70 0.404 600 1323 7.4 4.1 12.1 5.2 1.2 2.9 3 4 4
31.3 61.1 28.4 4.81 1.19 0.503 2.18 0.313 473 1201 5.8 3.7 11.2 4.8 1.1 2.4 4 3 3
Cerro Chachahuén DR8 DR9 intrusive agglom 67.70 60.56 0.36 0.83 14.91 17.27 3.27 5.38 0.16 0.14 0.94 1.42 3.85 5.53 4.61 4.62 4.13 3.70 0.14 0.29 100.07 99.75 26.6 52.1 15.9 3.9 1.06 0.434 1.77 0.280 548 1101 5.1 3.0 7.4 3.8 0.9 4.6 8 4 6
33.9 65.3 27.6 5.70 1.40 0.720 2.64 0.391 622 1071 3.4 3.4 10.7 5.2 1.1 7.6 <1 2 11
8.38 8.32 8.74 8.32 3.49 6.20 3.48 3.78 6.4 6.5 6.8 5.9 2.0 2.2 2.2 2.2 13.2 14.3 15.0 12.8 0.82 0.89 0.96 0.82 37.2 38.4 41.4 31.6 1060 1075 1263 990 28.5 28.0 30.5 31.3 0.34 0.36 0.28 0.31 0.24 0.23 0.23 0.21 2.32 2.34 1.97 2.06 4096 5690 6663 8983 2.9 3.0 2.5 3.2 36°59.21 36°59.41 37°4.28 37°4.33 68°52.55 68°52.11 68°51.45 68°51.45 1469 ± 42 1370 ± 62 1742 ± 39 1726 ± 57
Sierra de Chachahuén Arroyo Los Golpeados DRC6 DRC5b DRC5a welded tuff pumice clast 71.30 62.22 like5B 0.17 0.52 15.22 17.15 1.24 4.30 4.30 0.28 0.17 0.16 1.75 1.44 5.92 4.98 4.10 4.07 5.18 3.60 0.18 99.97 99.91 27.2 55.2 17.8 3.48 0.91 0.434 2.34 0.330 255 861 4.3 4.3 12.6 5.4 1.2 0.5 1 2 2
35.3 69.6 29.2 5.50 1.41 0.667 2.48 0.328 812 1045 4.6 3.7 11.5 4.9 1.3 6.5 10 5 8
33.8 67.6 25.2 5.15 1.30 0.627 2.52 0.344 621 815 7.9 3.4 10.9 4.5 1.1 6.3 9 4 8
10.16 7.75 7.8 1.5 11.6 0.88 31.7 717 22.6 0.46 0.22 2.35 10028 3.0 37°05.44 68°54,72 ~1320
7.70 2.46 6.4 2.2 14.2 0.88 29.6 815 27.5 0.33 0.26 2.35 6437 3.1 37°05.44 68°54,72 ~1320
7.70 2.46 6.6 2.0 13.4 0.86 24.1 729 30.2 0.32 0.25 2.41 3.2 37°05.44 68°54,72 ~1320
Cerro Corrales DRC10a DR28 pumice intrusive 63.28 61.94 0.57 0.53 18.51 17.77 4.07 4.84 0.14 0.19 1.26 1.59 5.11 5.20 4.08 4.55 2.78 3.41 0.18 0.26 99.98 100.27 34.6 76.8 31.7 5.89 1.58 0.694 2.64 0.344 942 1051 7.8 2.7 9.3 5.3 1.1 4.6 5 3 8
33.5 68.0 26.7 5.30 1.50 0.698 2.62 0.339 819 1038 4.3 3.1 11.6 4.8 1.1 6.1 7 8 9
6.86 7.96 3.24 3.03 5.9 6.3 2.2 2.0 13.1 12.8 0.92 0.94 30.4 31.0 929 907 30.6 29.3 0.27 0.35 0.21 0.24 1.77 2.43 2961 6591 3.5 3.7 37°04.35 37°4.21 68°54.56 68°55.31 ~1510 1698 ± 45
Cerro Bayo DR14 flow 62.60 0.49 17.88 4.08 0.27 0.95 5.53 4.68 3.25 0.22 99.95 33.4 68.1 32.7 6.51 1.72 0.793 2.80 0.432 1006 1301 4.7 4.1 9.8 4.5 0.9 3.3 4 5 5 7.94 4.28 5.1 2.3 11.9 0.90 39.0 1373 35.2 0.29 0.21 2.18 5798 2.4 37°2.98 68°51.57 1434 ± 37
52.66 1.05 18.45 8.45 0.22 2.47 9.91 4.24 2.07 0.53 100.05
31.4 67.9 38.1 7.53 2.16 0.989 2.97 0.406 1030 841 2.0 2.1 6.0 4.2 0.9 13.9 29 19 22
6.31 3.43 4.2 2.5 10.6 0.95 26.8 907 33.9 0.19 0.22 1.44 8419 2.9
SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O P2O5 Total
La Ce Nd Sm Eu Tb Yb Lu Sr Ba Cs U Th Hf Ta Sc Cr Ni Co
Na2O + K2O FeO/MgO La/Sm Sm/Yb La/Yb Eu/Eu* Ba/La Ba/Ta La/Ta Th/La Ta/Hf Th/Hf K/Cs Th/U Latitude °S Longitude °W Elevation (m)
Boina DR24a clast
Cerro
5.80 3.07 4.3 2.5 10.5 0.92 26.2 734 28.0 0.21 0.25 1.42 4533 2.8 37°04.35 68°54.24
25.0 52.3 29.0 5.86 1.62 0.762 2.38 0.331 836 654 4.0 1.9 5.2 3.6 0.9 15.0 39 19 21 6.20 4.05 3.7 2.6 9.9 0.88 30.6 1056 34.5 0.23 0.19 1.55 18686 2.9 37°04.35. 68°54.56 ~1510
24.1 53.7 29.1 6.43 1.72 0.871 2.44 0.336 779 738 1.0 1.9 5.6 3.6 0.7 15.1 5 7 18
base south side DRC11 DRC10b clast clast debris flow 54.98 51.48 1.13 1.24 18.01 18.66 7.92 8.45 0.13 0.16 2.58 2.08 8.58 9.86 3.62 3.90 2.18 2.30 0.42 0.45 99.56 98.58
6.26 2.64 4.1 2.5 10.3 0.93 30.8 916 29.8 0.22 0.27 1.77 1796 2.9 37°4.69 68°54.62 1403±35
29.2 62.2 29.7 7.20 2.05 0.991 2.83 0.393 1042 901 9.5 2.2 6.5 3.6 1.0 15.2 39 23 20
52.61 1.21 18.65 8.28 0.15 3.13 9.19 4.21 2.05 0.48 99.94
DR29 intrusive
6.56 3.05 4.3 2.4 10.4 0.96 27.9 814 29.2 0.21 0.22 1.37 8461 3.1 37°4.36 68°53.18 1450±30
28.1 59.5 29.8 6.46 1.92 0.914 2.70 0.404 858 784 2.5 1.9 5.9 4.3 1.0 13.5 17 13 21
55.09 1.35 18.35 7.86 0.20 2.58 7.60 3.97 2.59 0.40 100.00
DR5 lava
7.05 2.97 4.6 2.4 11.1 0.84 31.0 875 28.2 0.27 0.24 1.83 3672 2.9 37°4.36 68°53.18 1450±30
30.1 61.1 29.5 6.48 1.63 0.818 2.70 0.387 806 933 6.9 2.8 8.1 4.5 1.1 13.0 21 11 19 7.19 3.25 5.4 2.5 13.4 0.93 29.0 979 33.8 0.24 0.22 1.78 6041 2.6 37°4.36 68°53.18
32.4 62.5 30.0 6.01 1.65 0.743 2.42 0.328 719 940 4.2 3.1 7.9 4.4 1.0 9.6 17 12 15
DR6 DR7 clast clast block and ash 56.91 58.45 1.00 0.91 18.36 18.16 6.81 5.99 0.22 0.19 2.29 1.84 6.73 5.90 4.01 4.16 3.04 3.03 0.39 0.41 99.76 99.03
7.07 5.14 4.6 2.5 11.3 0.85 26.5 1054 39.7 0.20 0.19 1.52 8662 2.8 37°3.36 68°51.06 1403±20
37.0 79.1 41.6 8.01 1.97 0.936 3.27 0.452 972 980 2.3 2.6 7.4 4.8 0.9 6.7 3 4 11
57.21 0.81 19.00 6.95 0.21 1.35 6.88 4.62 2.45 0.39 99.89
south DR13 intrusive
7.05 5.09 5.4 2.4 12.6 0.87 29.9 884 29.5 0.29 0.24 2.01 5255 2.6 37°3.30 68°51.75 1480±37
39.4 81.9 40.8 7.35 1.82 0.81 3.11 0.444 920 1179 3.9 4.4 11.3 5.6 1.3 6.5 11 9 8
57.93 0.68 18.83 6.17 0.13 1.21 7.15 4.60 2.45 0.51 99.66
DR21 intrusive
6.22 2.24 4.9 2.4 11.8 0.88 24.0 948 39.5 0.17 0.20 1.33 5744 3.1 37°00.69 68°53.19 1499±28
34.5 68.6 33.7 7.11 1.87 0.921 2.93 0.408 913 828 2.9 1.9 5.8 4.3 0.87 11.4 41 25 17
56.03 0.91 18.13 6.96 0.23 3.10 7.25 4.21 2.01 0.35 99.19
DR33 intrusive
8.45 2.61 5.4 2.4 12.6 0.87 29.9 884 29.5 0.29 0.24 2.01 7254 2.6 37°01.64 68°53.70 1613±49
38.9 80.8 40.2 7.25 1.80 0.800 3.07 0.438 909 1164 3.8 4.4 11.2 5.6 1.3 6.4 11 8 8
58.38 0.77 17.10 4.41 0.13 1.69 8.46 5.12 3.34 0.25 99.65
DR31 dike
TABLE 3C. WHOLE-ROCK CHEMISTRY: CHACHAHUÉN VOLCANIC COMPLEX EARLY CHACHAHUÉN GROUP (MAFIC) Cerro Corrales Agrio Fm. Cerro "Condor" Cerro Bayo region Cerro Cabeza region
8.60 3.12 5.2 2.3 12.2 0.85 34.7 893 25.7 0.27 0.26 1.82 10291 2.5 37°01.2 68°53.5 ~1580
41.5 83.7 39.1 7.96 1.94 0.885 3.41 0.490 781 1441 2.8 4.6 11.3 6.2 1.6 6.8 5 5 10
56.72 0.98 16.93 5.41 0.26 1.74 9.07 5.10 3.50 0.36 100.07
DRC9a dike
5.81 2.40 4.8 2.6 12.6 0.88 33.4 868 26.0 0.25 0.29 1.88 2089 2.8 37°05.10 69°00.1
24.3 48.7 17.4 5.02 1.33 0.668 1.93 0.291 921 812 7.2 2.2 6.0 3.2 0.9 11.7 15 7 17
57.43 0.94 18.23 6.63 0.11 2.76 7.65 3.99 1.82 0.31 99.88
DRC24b flow
Cerro Vizcachas
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TABLE 3D. WHOLE-ROCK CHEMISTRY: LATE CHACHAHUÉN GROUP (SILICIC)
SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O P2O5 Total La Ce Nd Sm Eu Tb Yb Lu Sr Ba Cs U Th Hf Ta Sc Cr Ni Co Na2O + K2O FeO/MgO La/Sm Sm/Yb La/Yb Eu/Eu* Ba/La Ba/Ta La/Ta Th/La Ta/Hf Th/Hf K/Cs Th/U Latitude °S Longitude °W Elevation (m)
Group 1 Puesto Cerro Bayo Isaac west DR36 DR53 block and ash intrusive? dome 64.25 63.91 0.47 0.45 16.53 17.53 3.92 3.98 0.24 0.16 1.53 0.83 4.56 4.60 4.54 4.80 3.80 3.31 0.19 0.12 100.04 99.68
Complejo Volcanico DR30 block debris flow 58.08 0.66 17.65 5.25 0.32 1.64 7.49 5.44 3.29 0.18 100.00
Group 2 Cerro Corrales south DR25 Unit 6 pumice 60.97 0.54 17.21 4.49 0.22 2.22 6.93 3.87 3.38 0.17 99.99
Great Rodados Dike DR48 DR54 dike ignimbrite Unit 6 63.15 64.90 0.49 0.43 17.29 16.45 4.24 3.60 0.19 0.21 1.16 0.87 4.99 4.36 4.26 4.89 3.75 4.17 0.25 0.15 99.77 100.03
41.7 81.2 27.8 5.75 1.37 0.621 2.51 0.415 633 1329 10.8 4.7 13.1 6.4 1.5 4.8 3 3 7
42.6 83.0 31.5 6.41 1.64 0.820 3.07 0.473 755 1368 8.0 4.5 13.6 5.8 1.5 3.3 1 4 5
48.6 102.8 43.7 8.00 2.26 1.012 3.60 0.515 1240 1440 7.1 4.4 13.0 6.0 1.4 4.6 5 6 8
32.2 78.5 25.8 6.28 1.50 0.735 2.98 0.406 992 872 6.1 2.6 9.5 4.9 0.9 4.1 9 6 6
32.1 65.6 32.6 6.01 1.42 0.647 2.63 0.393 719 1147 4.3 3.8 9.7 4.8 0.9 3.9 0 2 6
32.3 65.6 23.4 5.34 1.47 0.627 2.64 0.401 805 1078 4.7 3.3 9.3 4.4 0.9 2.5 0 1 4
8.35 2.56 7.3 2.3 16.6 0.84 31.9 882 27.7 0.31 0.23 2.04 2919 2.8 37°1.74 68°54.52 1429 ± 59
8.11 4.82 6.6 2.1 13.9 0.86 32.1 937 29.2 0.32 0.25 2.35 3413 3.0 37°1.96 68°58.78
8.73 3.20 6.1 2.2 13.5 0.95 29.6 1051 35.5 0.27 0.23 2.15 3833 2.9 37°02.45 68°53.85 1711 ± 39
7.24 2.02 5.1 2.1 10.8 0.82 27.1 950 35.1 0.30 0.19 1.95 4602 3.6 36°04.1 68°55.3 ~1730
8.01 3.67 5.3 2.3 12.2 0.84 35.8 1274 35.6 0.30 0.19 2.04 7281 2.5 37°03.5 68°55.1 ~1950
9.06 4.15 6.1 2.0 12.2 0.95 33.4 1190 35.6 0.29 0.20 2.10 7313 2.8 37°0.44 68°56.84
Na2O + K2O FeO/MgO La/Sm Sm/Yb La/Yb Eu/Eu* Ba/La Ba/Ta La/Ta Th/La Ta/Hf Th/Hf K/Cs Th/U Latitude °S Longitude °W Elevation (m)
7.15 2.72 4.5 2.5 11.3 0.96 30.0 848 28.3 0.24 0.27 1.82 5203 2.5 37°2.79 68°55.14
6.48 2.28 4.4 2.3 10.3 0.93 31.4 914 29.2 0.29 0.23 1.91 3374 2.8 37°3.94 68°55.04 1694 ± 80
8.03 3.20 5.5 2.4 13.4 0.86 32.8 909 27.7 0.29 0.25 2.03 7244 2.6 37°3.67 68°55.33
8.68 3.34 5.3 2.2 11.6 0.82 30.6 1067 34.9 0.25 0.21 1.82 15784 2.7 37°1.74 68°54.52 1429 ± 59
38.4 77.0 39.0 7.31 1.79 0.930 3.32 0.491 928 1174 1.8 3.5 9.5 5.2 1.1 4.7 2 5 8 7.09 2.71 5.0 2.4 11.9 0.89 27.2 933 34.3 0.23 0.22 1.69 6276 2.7 37°03.3 68°52.3 ~1580
27.6 57.1 21.9 5.51 1.49 0.744 2.32 0.342 813 751 3.5 2.3 6.3 3.7 0.8 11.5 14 13 14 7.30 3.12 5.0 2.4 11.8 0.86 29.8 987 33.1 0.27 0.21 1.88 6943 2.7 37°06.1 68°53.2 ~1350
33.4 66.6 30.9 6.67 1.72 0.862 2.83 0.402 877 996 3.7 3.4 9.1 4.8 1.0 10.8 21 13 16 6.90 2.69 4.8 2.5 11.8 0.87 28.8 959 33.3 0.23 0.22 1.72 6490 2.6 37°3.82 68°54.99
30.1 61.8 26.5 6.27 1.67 0.862 2.55 0.385 934 867 3.5 2.7 6.9 4.0 0.9 12.9 16 12 16 6.70 2.26 4.9 2.3 11.2 0.91 33.2 1213 36.5 0.20 0.19 1.39 5572 2.8 37°4.42 68°51.42 1648 ± 39
23.6 49.3 23.5 4.83 1.32 0.626 2.11 0.317 720 784 4.0 1.7 4.8 3.5 0.6 10.6 6 7 14 6.28 2.31 4.9 2.7 13.1 0.94 29.2 1074 36.8 0.16 0.18 1.03 10740 2.6 37°4.55 68°51.18 1901 ± 37
34.3 69.4 34.8 6.96 1.96 0.891 2.61 0.367 947 1000 1.7 2.0 5.3 5.2 0.9 12.7 29 37 23
6.23 2.13 4.6 2.5 11.6 0.93 25.9 889 34.3 0.15 0.20 1.05 12308 2.8 37°4.55 68°51 1929 ± 20
26.3 54.8 23.0 5.73 1.65 0.807 2.27 0.322 853 681 1.4 1.4 4.0 3.8 0.8 17.5 29 18 24
6.15 2.25 3.8 2.8 10.7 0.89 31.0 918 29.7 0.17 0.21 1.05 5341 2.6 37°06. 68°51.24
24.1 53.8 22.4 6.40 1.75 0.896 2.25 0.298 900 745 3.2 1.6 4.1 3.9 0.8 15.6 31 35 26
La Ce Nd Sm Eu Tb Yb Lu Sr Ba Cs U Th Hf Ta Sc Cr Ni Co
34.0 66.6 28.9 6.21 1.60 0.798 2.54 0.388 895 1116 4.1 3.9 10.0 4.9 1.2 7.5 5 6 19
31.2 64.6 38.2 6.98 2.00 0.905 2.75 0.399 917 935 4.5 3.1 7.6 4.2 1.1 15.4 18 13 19
SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O P2O5 Total 27.3 58.2 27.0 6.18 1.75 0.840 2.65 0.382 879 857 6.7 2.8 7.8 4.1 0.9 15.3 15 13 21
TABLE 3E. WHOLE-ROCK CHEMISTRY: CHACHAHUÉN VOLCANIC COMPLEX LATE CHACHAHUÉN GROUP (MAFIC) Group 1 Group 2 Cerro Corrales region and Puesto Cerro Corrales region Cerro Chachahuén to the north - Unit 6 Isaac north side near top top south side DR47 DR26 DR46 DR37 DR24b DR49 DR45 DR12 DR11 DR10 DRC13 flow clast clast dike clast flow flow pyroclastic flow flow flow 55.32 55.39 56.94 57.65 57.04 58.60 57.73 59.19 56.06 52.71 53.82 1.01 1.10 0.74 0.61 0.85 0.85 0.82 0.68 0.72 1.25 1.35 17.99 18.30 18.82 18.38 18.56 17.90 18.19 17.48 18.17 18.20 17.53 7.17 7.20 5.48 5.16 6.51 6.42 6.28 5.63 7.21 8.16 8.67 0.14 0.17 0.20 0.19 0.15 0.21 0.17 0.15 0.22 0.20 0.17 2.64 3.15 1.71 1.55 2.40 2.06 2.33 2.49 3.12 3.84 3.85 7.40 7.67 6.97 7.11 7.00 6.31 7.33 6.66 6.92 9.06 8.06 4.30 3.77 4.49 5.17 4.43 4.22 4.14 4.01 4.11 4.13 4.10 2.85 2.71 3.54 3.52 2.66 3.08 2.76 2.69 2.18 2.10 2.05 0.37 0.27 0.45 0.28 0.24 0.26 0.27 0.29 0.41 0.50 0.49 99.18 99.72 99.34 99.61 99.84 99.91 100.03 99.27 99.11 100.16 100.10
Type
4.07 1.43 3.0 2.6 7.9 0.97 26.9 674 25.1 0.17 0.26 1.13 2867 3.1 37°3.26 68°52.35 1881 ± 88
16.6 38.1 20.1 5.45 1.68 0.828 2.10 0.273 794 445 3.6 0.9 2.9 2.5 0.7 33.1 522 170 50
DR22 clast 48.12 1.47 15.19 10.51 0.18 7.33 11.69 2.84 1.23 0.35 98.91
4.83 2.38 3.3 2.9 9.5 0.97 31.7 804 25.4 0.20 0.28 1.44 3611 2.9 37°2.53 68°58.95
21.5 47.2 32.7 6.48 2.00 0.940 2.26 0.327 954 680 3.0 1.5 4.4 3.0 0.8 24.3 47 24 32
DR52 flow 49.03 1.55 18.22 9.47 0.17 3.98 10.78 3.51 1.32 0.35 98.38
Cerro Boina
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TABLE 3F. WHOLE-ROCK CHEMISTRY: LATEST LATE CHACHAHUÉN GROUP AND AMPHIBOLE-RICH DIKES Latest Chachahuén Group (mafic) Amhibole-rich dikes Cañadón del Camino Cerro Montura Cerro Corrales - south side DR4 DR4N DR4B DRC9 DR39 DR40 DR27a DR27b DRC12 lava lava lava lava lava lava large ~east-west dike 48.32 48.32 50.11 53.38 53.67 SiO2 TiO2 1.82 1.82 1.29 0.85 0.92 Al2O3 16.98 16.98 19.19 19.09 19.15 FeO 11.00 11.00 11.00 8.96 7.70 8.04 12.98 8.81 9.24 MnO 0.20 0.20 0.20 0.26 0.25 MgO 4.78 4.78 3.30 2.40 2.35 CaO 11.53 11.53 9.64 9.18 8.89 3.52 3.49 3.38 4.00 4.00 4.02 1.06 2.09 2.93 Na2O K2O 1.26 1.26 1.74 2.23 2.25 P2O5 0.53 0.53 0.47 0.50 0.49 Total 99.94 99.90 98.89 99.59 100.03 La Ce Nd Sm Eu Tb Yb Lu Sr Ba Cs U Th Hf Ta Sc Cr Ni Co
25.3 55.3 30.1 7.15 2.03 1.03 2.78 0.376 984 550 5.1 1.6 5.4 3.4 0.6 29.2 26 20 37
24.6 56.0 34.8 7.87 2.17 1.11 2.84 0.403 1013 614 5.8 1.8 5.7 3.4 0.6 30.1 24 14 37
21.5 51.0 26.7 6.88 2.00 1.05 2.59 0.357 925 554 11.5 1.5 5.2 3.1 0.5 47.3 135 43 44
30.2 75.2 36.7 8.86 2.42 1.249 2.87 0.393 1173 770 5.9 1.9 7.3 3.3 0.7 16.5 3 3 28
31.1 68.7 31.4 7.59 2.01 0.945 2.91 0.434 1071 750 3.1 2.2 6.6 3.5 0.6 9.8 4 6 17
31.7 65.8 42.8 7.78 2.16 0.982 2.91 0.425 1056 827 1.2 2.3 6.3 3.5 0.6 10.2 10 8 17
Na2O + K2O 4.78 4.75 5.73 6.23 6.27 FeO/MgO 2.30 2.30 2.71 3.21 3.42 La/Sm 3.5 3.1 3.1 3.4 4.1 4.1 Sm/Yb 2.6 2.8 2.7 3.1 2.6 2.7 La/Yb 9.1 8.6 8.3 10.5 10.7 10.9 Eu/Eu* 0.91 0.89 0.92 0.88 0.89 0.93 Ba/La 21.7 25.0 25.8 25.5 24.1 26.1 Ba/Ta 938 1002 1017 1173 1236 1321 La/Ta 43.2 40.1 39.5 46.0 51.3 50.7 Th/La 0.22 0.23 0.24 0.24 0.21 0.20 Ta/Hf 0.17 0.18 0.17 0.20 0.18 0.18 Th/Hf 1.61 1.68 1.65 2.20 1.90 1.82 K/Cs 2068 1806 2441 5987 15823 Th/U 3.4 3.1 3.5 3.8 3.0 2.7 Latitude °S 37°6.1 37°0.68 37°0.75 37°6.08 37°6.08 37°6.08 Longitude °W 68°52.55 68°52.55 68°52.55 68°52.35 68°57.67 68°57.67 Elevation (m) 1300 ± 34 1300 ± 34 ~1320 1666 ± 48 1531 ± 31
5.0 2.6 7.9 3.54 0.33 0.447 1.46 0.206 214 173 0.2 0.1 0.1 0.9 0.1 68.1 38 40 53
18.6 41.4 19.5 5.11 1.57 0.821 2.06 0.262 758 503 1.9 0.9 3.0 2.3 0.9 17.2 15 13 29
29.3 66.4 32.8 7.74 2.26 1.06 2.55 0.359 875 705 1.4 2.3 6.1 3.5 0.6 20.4 25 25 31
1.4 2.4 3.4 0.31 34.6 2652 76.6 0.02 0.07 0.10 0 0.9 37°04.1 68°55.4 ~1710
3.6 2.5 9.0 0.95 27.0 546 20.2 0.16 0.40 1.31 0 3.3 37°04.1 68°55.4 ~1710
3.8 3.0 11.5 0.96 24.1 1266 52.6 0.21 0.16 1.77 0 2.7 37°04.35 68°54.24 ~1480
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Sample
TABLE 4. Sr, Nd, AND Pb ISOTOPIC RATIOS FOR CHACHAHUÉN VOLCANIC COMPLEX 87 143 206 All present-day ratios SiO2 Sr/86Sr Nd/144Nd Pb/204Pb 207Pb/204Pb Nd (wt%)
208
Pb/204Pb
Vizcachas Group DR44 Cerro de Las Vizcachas rhyolite DR20 Unit 2 dacite vitrophyre - south of Cerro Bayo DRC7C Unit 2 pumice
70.73 63.76 66.82
0.704449 0.704082 0.704008
0.512680 0.512734 0.512742
+0.8 +1.9 +2.0
18.550
15.593
38.366
Early Chachahuén Group DR5 Andesite flow top of Cerro “Condor” (“unit 5”)
55.09
0.703828
0.512784
+2.9
18.598
15.583
38.383
Late Chachahuén Group DR25 Cerro Corrales pumice (“unit 5”) DRC13 Long flow Cerro Chachahuén
60.97 53.82
0.703960 0.704148
0.512765 0.512743
+2.5 +2.1
18.628
15.609
38.476
53.67
0.703980
0.512758
+2.3
18.621
15.587
38.407
Latest Late Chachahuén Group DR40 La Montura basaltic andesite flow Note: Analytical methods are discussed in Appendix 3.
In a general way, the Early Chachahuén group can be divided into a more silicic, older part and a more mafic, younger part. The andesitic to dacitic rocks in the older part are largely related to explosive eruptions, domes, and shallow-level porphyries. They are notable for their amphibole phenocrysts and their complexly zoned plagioclase phenocrysts that contain glass inclusions. Clinopyroxene phenocrysts are common in the andesites. Analyses for dacitic samples (66–71% SiO2) in Table 3B are from shallow intrusive domes at Cerro de la Gloria (DR34 and DR35) and Cerro Chachahuén (DR8), and a welded tuff south of Cerro Corrales (DRC6). Analyses for andesitic samples (60–63% SiO2) are from a clast in an agglomerate in Cerro Chachahuén (DR9), a welded tuff (DRC6) and clasts in a pyroclastic flow (DRC5a and b) in the Arroyo los Golpeados south of Cerro Corrales, pumice from the base of Cerro Corrales (DRC10a), a dome southwest of Cerro Corrales (DR28), and the hornblende-bearing andesite in Cerro Bayo (DR14). The mafic part consists mainly of basaltic to mafic andesitic dikes and lava flows that cut and cap the silicic sequences. Analyses for mafic andesites in Table 3C (55–58% SiO2) are from agglomerate blocks, lava flows, and dikes in Cerro Vizcachas (DRC24b), near Cerro Cabeza (DR31, DRC9a, DR33), south of Cerro Bayo (DR21 and DR13), and from Cerro “Condor” (DR5, DR6, DR7). Those for basaltic andesites (52.5–55% SiO2) are from agglomerates at Cerro Corrales (DRC11, DRC10b) and Cerro Boina (DR24a), and a dike cutting the Cretaceous Agrio Formation south of Cerro Corrales (DR29). Chemical analyses of Early Chachahuén group samples are listed in Tables 3B and 3C and plotted in Figures 6A and 6C. Samples with ~52%–68% SiO2 plot in the high-K field on the K 2O-SiO2 diagram (Fig. 5A), in the trachyandesite, trachyte, and rhyolite fields on an alkali-SiO2 diagram, and mostly in the tholeiitic field on the SiO2 versus FeO/MgO Miyashiro plot (Fig. 5B). The more silicic samples (60%–71% SiO2) are like Vizcachas group samples in having high Ba (815–1200) and Sr (250–1000) concentrations and small negative Eu anomalies, but differ in having flatter REE patterns (La/Yb = 12–15) with lower La/Sm (6.5–7.8) and Sm/Yb (most 2–2.2) ratios (Fig. 7).
More mafic samples (52%–58% SiO2) are also Ba (740–1440) and Sr (720–1042) rich. Compared to the silicic samples, they have flatter overall REE patterns (La/Yb = 10–13), with lower light (La/Sm = 4.6–5.9) and similar heavy (Sm/Yb = 2.3–2.6) REE ratios (Fig. 6C). All Early Chachahuén group samples have higher La/Ta (23–40) and lower Ta/Hf (0.19–0.26) ratios than Vizcachas group samples (Figs. 8 and 9). These ratios reach arc-like values (La/Ta > 25; Ta/Hf < 2). A mafic andesite (DR5, Table 3) has lower 87Sr/ 86Sr (0.7038) and higher 206Pb/ 204Pb (18.60) ratios and a higher ε Nd (+2.9) values than the Vizcachas samples (Table 4; Fig. 10). Late Chachahuén Group The Late Chachahuén group includes the Complejo Volcanico, most of unit 5, all of units 6 and 7, and the Cerro Chachahuén flows, along with related dikes and intrusives (Table 1; Figs. 3C and 4B). In general, there are two sequences of andesitic to dacitic ignimbrite-pyroclastic units that are each overlain by basaltic to mafic andesitic lava flows and agglomerate units. The Complejo Volcanico consists of volcanic debrisflow material, which, along with unit 5 pyroclastic flows and agglomerates, occurs on the east side of a fault zone running from Cerro Corrales to the north on the Pérez and Condat (1996) map. Unit 6 includes the thick sequences of tuffs and tuffaceous “mass flows” interbedded with volcanic agglomerates that overlie and occur west of the Complejo Volcano and unit 5. Unit 6 is topped by gray columnar jointed lava flows. Unit 7 consists of the basal sandy conglomerate overlain by columnar jointed lava flows and thick pyroclastic deposits that form the upper part of Cerro Corrales. The Late Chachahuén group comprises the prominent sequence of 4–6-m-thick mafic andesitic lava flows interbedded with thin pyroclastic layers at the Cerro de Chachahuén. Also included in this group is the andesitic “great dike,” which is up to 45 m wide and runs northwest from Cerro Corrales (Fig. 3C), and dikes up to 6 m wide on the south side of Cerro Corrales. The latter can be distinctive in having very large hornblende crystals (up to 15 cm across) with inclusions of clinopyroxene and titanomagnetite set in a fine-grained groundmass.
Evolution of the late Miocene Chachahuén volcanic complex
Figure 4. Thematic mapper images corresponding to general region in maps in Figures 3A, 3B and 3C showing proposed eruptive centers (dashed ovals) and faults (dashed lines). Eruptive center labeled A at Cerro de las Vizcachas is not shown in Figure 3.
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K-Ar ages reported by Pérez and Condat (1996) for the Late Chachahuén group as constituted here range from 6.4 ± 0.4 to 5.6 ± 0.3 Ma. A new groundmass 40Ar/ 39Ar age of 4.85 ± 0.03 Ma (DRC13 in Table 2) on a mafic andesite lava flow on the southwest side of Cerro Chachahuén (Fig. 3C) extends the younger age limit into the early Pliocene. Chemical analyses of Late Chachahuén group samples are listed in Tables 3D to 3E and plotted in Figures 6B and 6C. The samples can generally be put into chemically coherent silicic and mafic subgroups (Fig. 6B). One silicic subgroup consists of dacitic samples from a block and ash deposit at Puesto Isaac (DR36) and a dome west of Cerro Bayo Chico (DR53). A second includes andesitic to dacitic samples from a “unit 6” pumice in Cerro Corrales (DR25), the “great dike” (DR48), and an ignimbrite near Cerro Rodados (DR54). An andesitic block in the Complejo Volcanico (DR30) is also similar. A first mafic subgroup includes mafic andesitic samples (55%–57% SiO2) from clasts (DR46, DR26) in agglomerates and a flow (DR47) on or north of Cerro Corrales. A second includes samples from a dike cutting the Complejo Volcanico near Puesto Isaac (DR37), and a block from a block and ash deposit (DR24b) and lava flows (DR49, DR45) on and north of Cerro Corrales. Other mafic samples (53%–59% SiO2) are from Cerro Chachahuén flows and pyroclastic rocks (DR10 to DR13). Basaltic samples (48%–49% SiO2) are from a lava flow west of Cerro Campanarío (DR52) and an agglomerate near a NW-trending fault at Cerro Boina (DR22). Analyses of the amphibole-bearing dikes (DR12, DR27a, and DR27b) on the south side of Cerro Corrales are included in Table 3F. To a large degree, Late Chachahuén group samples are chemically similar to Early Chachahuén group samples (Figs. 5–10). Most plot in the high-K field in Figure 5A, in the trachyandesite field on an alkali-SiO2 diagram, and in the tholeiitic field in Figure 5B. Those with 60%–65% SiO2 have similar Ba (872–1368 ppm) and Sr (633–1240 ppm) concentrations, small negative Eu anomalies, and REE patterns (La/Yb = 11–17; La/Sm = 5.1–7.3; Sm/Yb = 2–2.3) to Early Chachahuén samples with similar SiO2 contents. The same analogy is true for samples with 52–59% SiO2 (681–1174 ppm Ba; 713–934 ppm Sr; La/Yb = 10–13; La/Sm = 3.8–5.0; Sm/Yb = 2.0–2.8). Other similarities with Early Chachahuén samples include ranges of La/Ta (28–37), Ta/Hf (0.18–0.27), and Ba/La (26–36) ratios (Figs. 7 and 8). As seen in Figure 10 and Table 4, two Late Chachahuén andesites with 54% and 61% SiO2 have higher 87Sr/ 86Sr (~0.7040) and 206Pb/ 204Pb (18.63) ratios and lower εNd (+2.1–+2.5) values than an Early Chachahuén andesite with 55% SiO2. Two Late Chachahuén group basalts (48–49% SiO2) differ from less mafic samples in having flatter REE patterns (La/Yb = 8–10) with lower La/Sm (3–3.3) and somewhat higher heavy Sm/Yb (2.6–2.9) ratios, and slightly lower La/Ta (~25) and higher Ta/Hf (0.26–0.28) ratios (Figs. 6B, 6C, 7, and 8). The Cerro Boina basalt (DR22) is notable for its primitive characteristics (>500 ppm Cr; 170 ppm Ni; >7% MgO).
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Figure 5. (A) K2O (wt%) vs. SiO2 (wt%) for late Miocene to earliest Pliocene Chachahuén volcanic complex samples (points) on arc volcanic rock classification diagram. Also shown are fields for 24–20 Ma and younger than 3 Ma alkaline volcanic rocks from the Sierra de Chachahuén region and for latest Miocene to earliest Pliocene backarc Pocho and San Luis volcanic rocks from over the Chilean flat-slab to the north (Fig. 1). Data for Chachahuén volcanic complex samples are from Tables 3A to 3F, for alkaline volcanic rocks are from Kay et al. (this volume, chapter 2), and for Pocho and San Luis volcanic rocks are from Kay and Gordillo (1994). (B) Plot of same samples as in A on FeO/MgO vs. SiO2 (wt%) classification diagram for arc tholeiitic and calc-alkaline arc volcanic rocks from Miyashiro (1974).
Latest Late Chachahuén Group The Late Chachahuén group also includes basaltic to basaltic andesite flows on the periphery of the Chachahuén volcanic complex (Table 3F; Figs. 6B and 6C). Analyzed samples are from basaltic andesite flows (DR39 and DR40) at Cerro Montura along a NW-trending fault zone in the northwest part of the Sierra de Chachahuén, and the Cañadón del Camino basalt flow (DR4 and 4N, DR4b, DRC9) in the southern part of the range. These flows
are clinopyroxene-rich basalts and basaltic andesites (50%–54% SiO2) that have REE patterns (La/Yb = 8–11) marked by moderate light (La/Sm = 3.1–4.1) and heavy REE slopes (Sm/Yb = 2.6–3.1) and Ba/La ratios from 21 to 26 (Figs. 6 and 7). They are notable for their higher La/Ta (40–51) and lower Ta/Hf (0.17–0.20) ratios compared to other Late Chachahuén samples (Figs. 8 and 9). The isotopic ratios of a Cerro Montura basaltic andesite are like those of a Cerro Chachahuén basaltic andesite flow (Fig. 10).
Figure 6. Extended trace-element plots (A and B) and rare earth element (REE) plot (C) for Chachahuén volcanic complex samples. Primitive mantle mormalization factors (all in ppm) for extended trace-element plots are from Sun and McDonough (1989): Cs (0.032), Ba (6.989), Th (0.085), U (0.021), K (250), Ta (0.041), La (0.687), Ce (1.775), Sr (21.1), Nd (1.354), Sm (0.444), Hf (0.309), Eu (0.168), Tb (0.108), Yb (0.493), and Lu (0.074). Normalization factors (all in ppm) for REE plots are from the Leedey chondrite: La (0.378), Ce (0.976), Nd (0.716), Sm (0.23), Eu (0.0866), Tb (0.0589), Yb (0.249), Lu (0.0387), Th (0.05), Ta (0.02), and Sr (116).
Evolution of the late Miocene Chachahuén volcanic complex 231
Figure 6 (continued).
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Evolution of the late Miocene Chachahuén volcanic complex
Figure 6 (continued).
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Figure 7. (A) La/Yb ratio vs. SiO2 (wt%), and (B) La/Sm vs. Sm/Yb ratios for same samples as in Figure 5. La/Sm and Sm/Yb ratios indicate the slopes of the light and heavy rare earth element (REE) patterns, respectively. See text for discussion.
Figure 8. (A) La/Ta ratio vs. SiO2 (wt%), and (B) Ba/Ta vs. La/Ta ratios for same samples as in Figure 5. High La/Ta, Ba/Ta, and Ba/La ratios are indicators of components associated with the subducting plate and overlying mantle wedge.
DISCUSSION: ERUPTION HISTORY, MAGMA EVOLUTION, AND TECTONIC SETTING OF THE CHACHAHUÉN VOLCANIC COMPLEX
Early Eruptive Centers (Vizcachas and Chachahuén Groups) The eruption of the Chachahuén volcanic complex began with the silicic andesitic and dacitic units of the Vizcachas group. One center (marked A in Fig. 4A) is needed to explain the distribution of Vizcachas group outcrops in the Cerro de Las Vizcachas, and others are needed to explain the broad NE-SW– trending band of Vizcachas group outcrops crossing the Sierra de Chachahuén (Fig. 3A). A series of curved faults associated with these outcrops points to possible centers in, southwest of, and north of the Cerro Bayo region. The overlap of the band of Vizcachas group outcrops with exposures of the Cretaceous Agrio Formation and early Miocene basalt flows, which are considered to be associated with a NE-trending fault system (Fig. 3A), is consistent with localized Vizcachas group vents along the same fault system. The concentration of silicic vitrophyres in the Cerro Bayo region can be reconciled with rapid
Eruptive History of the Chachahuén Volcanic Complex Chachahuén complex volcanic units can be linked to eruptive centers based on their distribution, geomorphic expression, chemistry, and ages. The overall picture is one of a series of eruptions along faults and from small nested downsag calderas (e.g., Lipman et al., 1996). The younger centers are easier to identify, because the older ones are more eroded and partially covered by volcanic debris flows, alluvial, and colluvial deposits, and Pliocene to Quaternary alkali basaltic lava flows. The generalized eruption scheme and possible eruptive centers shown in Figures 3 and 4 are discussed in the following sections.
Evolution of the late Miocene Chachahuén volcanic complex
Figure 9. Th/Hf vs. Ta/Hf ratios for same samples as in Figure 5. Th/Hf ratios are relative measures of source components associated with a subducting oceanic slab, and Ta/Hf ratios are relative measures of mantle enrichment. General ranges of values observed in intraplate, backarc, and arc are indicated. MORB—mid-ocean-ridge basalt.
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eruption of superheated magmas at the intersection of the NEtrending fault system with a NW-trending system (Figs. 3A and 4A). The chemistry of Cerro de Las Vizcachas and Cerro Bayo region dacites shows that similar composition magmas erupted from diverse centers (Table 3A). The ca. 6.9 Ma to 6.3 Ma Early Chachahuén group marks an increase in volcanic activity in the same general region. Candidates for eruptive centers are marked by the letters B, D, and E on the TM image in Figure 4A. A problematical center at C, marked by a prominent circular feature near Cerro Corrales, is in a region covered by alluvial and colluvial deposits and basalt flows. The prominent oval (B) just west of Puesto Zuñiga is a candidate for a highly dissected vent for both the Vizcachas and Early Chachahuén Groups. The northern rim of B is marked by a fault between the Cretaceous Agrio Formation and the Vizcachas group. The Vizcachas stage is dated by K-Ar ages of 7.3 ± 0.4 Ma on a dacite dome (?) along a fault, and 7.28 ± 0.07 Ma on a clast
Figure 10. (A) 87Sr/ 86Sr ratios vs. SiO2 (wt%) and εNd values for Chachahuén volcanic complex samples relative to fields for early Miocene and younger than 3 Ma Sierra de Chachahuén region alkaline volcanic rocks and Pocho and San Luis volcanic rocks from over the Chilean flat-slab to the north. SVZ—Southern Volcanic Zone. (B) 208Pb/ 204Pb and 207Pb/ 204Pb ratios vs. 206Pb/ 204Pb ratios for same samples as in A. Data for Chachahuén volcanic complex are from Table 4; other data are from references in Figure 5. NHRL—Northern Hemisphere reference line.
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in a pyroclastic flow east of Puesto Zuñiga. The late Vizcachas and Early Chachahuén stages at centers marked B and C are dated by K-Ar ages of 7.1 ± 0.4 Ma on an ignimbrite near the northwest rim, of 7.0 ± 0.4 Ma and 6.8 ± 0.3 Ma on andesite intrusives into Mesozoic units, and of 6.1 ± 1.3 Ma on a unit 3 andesite, 6.4 ± 0.3 Ma on an intrusive, 6.5 ± 0.3 Ma on a dike, and 6.5 ± 0.3 Ma on a flow in the Arroyo los Golpeados region to the south. Another possible vent (C) in the Cerro de La Pascua region north of Cerro Bayo is largely obscured by younger basalt flows. The margins are marked by a fault along the northwestern margin, a curved anticline axis to the south, and pyroclastic deposits. The Cerro La Pascua intrusive could be a dome on the rim. Activity in this region is dated by K-Ar ages of 7.0 ± 0.4 Ma on a basaltic andesite lava flow along the circular fault in the north and of 6.5 ± 0.4 Ma on a unit 3 andesite lava flow near the northern margin (Figs. 3A and 3B). A candidate (D) for a complex center important in the Early Chachahuén stage is in the region of Cerro Bayo and Cerro Chachahuén. The southern part is partially obscured by the Late Chachahuén group. Early Chachahuén stage activity in this region is dated by K-Ar ages of 6.9 ± 0.3 Ma on a flow capping Cerro Raton, 6.7 ± 0.3 Ma on the Bayo dome, 6.6 ± 0.3 Ma on a basaltic andesite intrusive near Puesto J.A. Sosa, 6.4 ± 0.3 Ma on an andesite flow on the north side of Cerro Chachahuén, and 6.3 ± 0.3 Ma on an andesite south of Cerro Boina. Central Eruptive Centers (Late Chachahuén Group) Vents and structures related to the Late Chachahuén group are somewhat clearer. The origin of this group is best attributed to a central caldera with satellite centers. A principal candidate for a caldera is the distorted ring structure in the central part of the Sierra de Chachahuén (Figs. 3C and 4). A segment of the western wall is marked by the boundary fault separating unit 6 from the Complejo Volcanico and unit 5 pyroclastic flows in the interior of the caldera. Two sequences erupted in this region. The first includes unit 5 andesitic to dacitic ignimbrites and associated pyroclastic deposits at the base and east of the wall, and east of Cerro Corrales to the south. The second includes the overlying unit 6 tuffs and mafic andesite flows that appear to have emanated from a fracture defining the caldera rim. The Complejo Volcanico, east of the wall, represents a volcanic collapse (debris flow) into the interior of the caldera. Chaotic deposits in the Cerro Corrales region are probably hot avalanche or debris-flow deposits. Dikes that radiate into the center (including the “great dike”) cut these sequences. The mafic andesites in Table 3E form chemically coherent groups, which can be associated with the first (DR37, DR49, DR24b, DR45) and second (DR47, DR26, DR46) eruptive events. Chemical similarities of ignimbrites (Table 3D) from Cerro Corrales (DR25) and near Cerro Rodados (DR54), the Complejo Volcanico (DR30) collapse, and the “great dike” (DR48) are in accord with all of these samples being related to the younger eruptive cycle. The Complejo Volcano debris flow is also associated with this cycle.
The irregular shape of the caldera and the distribution of the erupted material can be explained by a trap-door caldera. Lavas from this type of caldera erupt through an uplifted trapdoor fault as other margins of the structure collapse along faults. The faulted margins of the caldera (labeled 1–3; Fig. 3C) are: (1) a NE-SW–trending fault running from the Río del Cafe toward Cerro Chachahuén, (2) a SE-trending fault that passes along the early Miocene outcrops northwest of Cerro Chachahuén, and (3) a NE-SW–trending fault system associated with the early Miocene and Vizcachas group outcrops along the Arroyo los Golpeados. The occurrence of a basaltic flow (DR22) in Cerro Boina near the intersection of faulted margins 2 and 3 is consistent with ascent of primitive magmas at the junction of faults. Both the dike orientations, which generally parallel the boundary fault segments, and the NE- and SWtrending fault systems could be controlled by the same stress systems that produced these faulted margins. Eruptive units that have K-Ar ages that are consistent with the “trap-door” caldera forming at ca. 6.3 ± 0.3 Ma are found as caldera-related deposits at the base of unit 6 at Cerro Corrales and in the southeastern part of the caldera. A hornblendebearing dike cutting these deposits in Cerro Corrales has a K-Ar age of 6.1 ± 0.3 Ma. K-Ar ages of 5.9 ± 0.3 Ma that come from two andesitic outcrops west of the Cañadón del Camino in the Arroyo los Golpeados are interpreted as parts of a debris flow. The most obvious satellite center is marked by the Cerro Chachahuén flows at Cerro Chachahuén (Figs. 3C and 4). The remains of a similar center in the Cerro de los Rodados region are suggested by remnant lava flows on the TM scene (Fig. 4). These centers along with smaller possible vents south of the Río Café are localized near the prominent NW-SE fault system that forms part of the northern boundary of the trap-door caldera. There is a spatial and temporal overlap of these satellite centers with flows along the main caldera rim. K-Ar ages are 6.3 ± 0.3 Ma for a flow near the top of the Cerro Chachahuén center and 6.4 ± 0.3 Ma for a pyroxene andesite flow near Cerro Tanque in the los Rodados center. A flow on the southwest side of Cerro Chachahuén yielded a 40Ar/ 39Ar age of 4.85 ± 0.03 Ma (Table 2). Post–Central Caldera (Late Chachahuén Group) The last events in the Chachahuén volcanic complex postdate caldera formation. Collapse of a dome near Puesto Isaac could explain the destruction of the central part of the los Rodados center. Nearby outcrops mapped as “intrusives” could be domes, dacitic ignimbrites, and block and ash deposits produced by dome explosion or collapse. Such an event explains chemical similarities (Table 3D) between a dacite (DR53) to the west and a block and ash deposit to the north of Puesto Isaac (DR36). A debris flow would explain the hummocky valley fill on the TM scene (Fig. 4). The time of dome collapse could be constrained by a K-Ar age of 6.1 ± 0.3 Ma on an andesite north of Puesto Isaac. Similar shallow-level intrusive domes formed near this time or shortly after comprise the andesitic-dacitic
Evolution of the late Miocene Chachahuén volcanic complex “intrusives” below the Chachahuén lava flows at the Cerro Chachahuén center. Their age is constrained by K-Ar dates of 6.0 ± 0.3 Ma at Cerro Bombilla and 5.6 ± 0.3 Ma on the northern side of Cerro Chachahuén. The last events seem to have been eruptions of peripheral lava flows like those at Cerro Montura and in the Cañadón del Camino. These flows could be similar in age to the 4.85 ± 0.03 Ma flow on the southwestern side of Cerro Chachahuén. Origin and Evolution of the Magmas of the Chachahuén Volcanic Complex The distinctive petrologic, geochemical, and isotopic signatures of the Chachahuén volcanic complex rocks are important in understanding the origin and evolution of the Chachahuén volcanic complex. The chemical and isotopic data show that the magmas evolved from (1) a dacitic to rhyodacitic Vizcachas group with intraplate-like chemical tendencies to (2) basaltic to dacitic Early and Late Chachahuén groups with arclike tendencies that peaked in the Late Chachahuén mafic peripheral flows. These data are consistent with a model in which Chachahuén volcanic complex magmas contain crustaland mantle-derived components that have characteristics that have changed through time due to the increasing influence of subducted components above a shallowing subduction zone. Andesitic to Rhyodacitic Magmas Differences between major- and trace-element characteristics, phenocryst assemblages, and isotopic ratios are important in understanding the origin and evolution of Vizcachas group and Early to Late Chachahuén group silicic andesitic to rhyodacitic magmas. Evidence for source region differences comes from relatively lower K contents (Fig. 5), lower Th/U ratios, and intraplate-like rather than arc-like La/Ta (Fig. 8) and Ta/Hf (Fig. 9) ratios in the Vizcachas volcanic rocks. Other evidence comes from higher 87Sr/ 86Sr and lower εNd and 206Pb/ 204Pb ratios (Table 4; Figs. 10A and 10B) in Vizcachas relative to Early Chachahuén volcanic rocks. Greater differences in rhyolitic than in silicic andesitic rocks show that less different mafic magmas evolved into more distinctive silicic magmas. A role for crustal contaminants in the silicic magmas is consistent with low K/Cs ratios (<10,000) and high Cs contents (up to 8.5 ppm) and a tendency for 87Sr/ 86Sr ratios to increase and εNd to decrease with increasing SiO2 (Fig. 10A) within the Vizcachas group. Striking features of both Vizcachas and Early to Late Chachahuén group silicic rocks are their high Ba, Sr, and U contents. Particularly notable are Sr and U spikes in mantlenormalized trace-element patterns (Figs. 6A and 6B) and the contrast between high arc-like Ba/La ratios (32–46) and intraplate-like La/Ta (10–22) and Ta/Hf (>0.4) ratios in the Vizcachas group (Fig. 8). Alternatives for explaining high Ba, Sr, and U concentrations are that they reflect preexisting continental lithospheric mantle and/or crustal compositions or that they were introduced near the time of magma formation. Small
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negative Eu anomalies in the Vizcachas volcanic rocks (Fig. 6C) and a correlation of high Sr and Ba with high U rule out an important role for melting of feldspar-rich crust. Understanding the differences between the Vizcachas and Chachahuén silicic volcanic rocks requires examining possible contaminants from the underlying crust. A candidate for contributing to the intraplate-like character of Vizcachas group magmas is the late Miocene crustal basement. In this regard, the last major event affecting the deep crust of this region produced the extensive Choiyoi Province rhyolites, which are associated with the initial Mesozoic rifting of Pangea (e.g., Kay et al., 1989). This event left a depleted crustal residue that was complementary to the Choiyoi rhyolites, along with the crystallized remains of the intraplate mafic magmas that triggered melting of the crust (Pankhurst and Rapela, 1995). Deep crustal magmatic additions in the early Miocene related to alkaline volcanism (Kay and Copeland, this volume, chapter 9) would have enhanced the intraplate-like character of this crust. A lack of early Miocene silicic magmas attests to a relatively refractory melt-depleted deep crust in the early Miocene. This intraplatelike crust would be the crust to contribute a crustal contaminant to the Vizcachas magmas. The continued introduction of late Miocene mantle-derived magmas from above a subducting slab into the crust would progressively lead to a more arc-like crust, which would subsequently be the contaminant in the Early to Late Chachahuén group magmas. Evidence for differences in crystallization conditions between Vizcachas group and Early to Late Chachahuén group magmas comes from their phenocryst assemblages. A notable difference between them is that orthopyroxene and biotite phenocrysts occur in silicic andesites and dacites of the Vizcachas group, whereas amphibole and salitic clinopyroxene phenocrysts occur in andesites and dacites of the Early and Late Chachahuén groups. Common phenocrysts in all of these rocks are plagioclase, titanomagnetite, accessory apatite, and sometimes titanite. The presence of magnetite and titanite highlights the oxidizing conditions under which all of the magmas evolved. Complex zoning and glass inclusions in pyroxene, amphibole and plagioclase phenocrysts, amphibole rims on pyroxenes, and partially melted cores in plagioclase phenocrysts support some mixing of mafic and silicic magmas. Additional information on fractionating and residual mineral assemblages in the Vizcachas group and Early to Late Chachahuén group magmas comes from trace-element trends with increasing silica content. A pattern of decreasing REE concentrations in Vizcachas group andesites to rhyodacites that reaches an extreme in the middle and heavy REEs (Figs. 5A and 5B) requires removal of middle and heavy REE-bearing mineral phases. As argued by Luhr et al. (1984) and Kay and Gordillo (1994), the best candidates are apatite, titanite, and amphibole. A role for zircon is ruled out by high Hf, Th, and U concentrations that increase with Si content. Small Eu anomalies (Eu/Eu* = 0.90–1.06) and very high Sr (640–1040 ppm) and Ba (>900 ppm) contents restrict removal of plagioclase and
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K-feldspar. In contrast, Early and Late Chachahuén group silicic andesitic to rhyodacitic lavas exhibit less heavy and middle REE depletion and more concave-up REE patterns (Fig. 6), which is consistent with a lesser role for titanite and a greater role for amphibole. Larger negative Eu anomalies (Eu/Eu* = 0.82–0.94) and lower Sr contents (as low as 250–600) at similar Ba contents are indicative of a more important role for plagioclase fractionation or residual plagioclase in a crustal source region or both. As in the Miocene Pocho volcanic rocks over the Chilean flat-slab near 32°S latitude (Fig. 1; Kay and Gordillo, 1994), trace-element patterns and phenocryst assemblages of the Vizcachas group volcanic rocks can be explained by removal of orthopyroxene, titanite, hornblende, magnetite, and limited plagioclase. Comparisons with crystallization experiments on Mexican arc andesites (63% SiO2) in Moore and Carmichael (1998) show that generally similar oxidized andesitic magmas with orthopyroxene followed by plagioclase and magnetite on the liquidus have ~5% H2O and begin to crystallize at temperatures near 950 °C and PH2O near 0.95 Mpa. At higher water contents (>5.5%) and pressures (>1.05 Mpa), hornblende is the liquidus phase followed by magnetite and plagioclase. Rapid eruption of such magmas from depths of ~6 km at temperatures above 900 °C would cause devolatilization and provide an explanation for the Vizcachas group vitrophyres. Pooling of the Vizcachas group magmas at depth before eruption is consistent with formation in a contractional regime. The localization of vents along faults is consistent with an association between Vizcachas group eruptions and the faulting that uplifted the Sierra de Chachahuén. These conditions are also consistent with the rhyodacitic Vizcachas group magmas being dominantly composed of an intraplate-like lower-crustal component derived from melting by and mixing with wet, fluid-mobile element-rich (U, Sr, and Ba) mantle magmas generated above a shallowing Nazca plate. The influx of hydrous fluids and melts can be argued to lead to thermal weakening and melting of the continental lithosphere and crust, facilitating contractional deformation (see Kay and Gordillo, 1994; James and Sacks, 1999; Ramos et al., 2002). Increasing proportions of newly introduced mantle melt in Vizcachas group silicic andesites and dacites can explain their transitional arc-like trace-element tendencies (Figs. 5–8). In contrast to Vizcachas group samples, phenocryst assemblages and trace-element characteristics of Early to Late Chachahuén group andesitic to rhyodacitic samples show evidence for crystallization of hornblende, plagioclase, salitic augite, and magnetite. Their trace-element trends indicate more plagioclase and less titanite fractionation (Fig. 6A–C), and their FeO/MgO ratios (Fig. 5B) suggest less magnetite fractionation than for the Vizcachas group magmas. A comparison with the experimental work in Moore and Carmichael (1998) shows that generally similar andesitic magmas that crystallize amphibole followed by plagioclase and augite on the liquidus have >6% H2O, an oxygen fugacity ~2 units above NNO, and PH2O > 1.09 Mpa. Eruption temperatures of <900 °C in the Early to Late Chachahuén
andesites are consistent with crystallization of augite after plagioclase. A lower eruption temperature than for the Vizcachas group magmas fits with the lack of vitrophyres. The Early to Late Chachahuén group phenocryst assemblages and the very large amphibole crystals (>20 cm) that occur in some of the dikes require eruption from depth because amphibole is unstable at <6 km. Pooling of Early to Late Chachahuén group silicic magmas at these depths is also consistent with slow ascent of hydrous magmas through a crust under contraction. Eruption of silicic magmas followed by basaltic to mafic andesitic magmas fits with a temporal relation between eruption of pooled fractionating magmas, the arrival of batches of mafic magmas from depth, and eruption associated with active faulting. Mineralogical and textural evidence for mixing supports eruptions linked to intrusion of mafic magmas into differentiating magma chambers. Mafic Mantle-Derived Magma The arc-like chemical signatures of the Early and Late Chachahuén group magmatic rocks can be linked to mantlederived magmas by the trace-element and isotopic signatures of the most mafic rocks. Evidence for subduction-related components in the mantle source comes from comparisons with arc magmatic rocks worldwide, which have ratios of La/Ta > 25, Ba/La > 20, and Ta/Hf < 1.5, and with Holocene Southern Volcanic Zone arc and near backarc magmatic rocks, which have La/Ta ratios of ~35–80 (see Hickey et al., 1986). For comparison, Early and Late Chachahuén basaltic to mafic andesitic lavas have ratios of La/Ta ~ 22–40, Ba/La ~ 25–40, and Ta/Hf ~ 0.19–0.30 (Figs. 7 and 8). Overlapping isotopic ratios in Early and Late Chachahuén group and Southern Volcanic Zone arc rocks are consistent with subducted sediment or crust removed by forearc erosion or both having been incorporated into the mantle source (e.g., Stern, 1991; Kay et al., 2005). Trace-element and isotopic data from the most mafic Early and Late Chachahuén group rocks support a progressive increase in the influence of subducted components during the late Miocene that can be associated with the cumulative influence of a hydrous, oxidized fluid component from a subducting oceanic slab and subducted crustal material. Importantly, the most arc-like ratios occur in the Late Chachahuén group basaltic to high-K andesitic peripheral flows (49–56% SiO2; La/Ta = 40–52; Ta/Hf = 0.17–0.20; Ba/La = 22–26) at Cañadón del Camino and Cerro Montura (Fig. 3C). More evidence for a temporal increase in the influence of subducted components comes from an Early Chachahuén mafic andesite (55% SiO2), which has a lower 87Sr/ 86Sr ratio and a higher εNd value than Late Chachahuén lavas with <54% SiO2 (Fig. 10A). Even lower 87Sr/ 86Sr and 206Pb/ 204Pb ratios at a given ε Nd in Vizcachas group volcanic rocks support the hypothesis that temporal changes in crustal and mantle sources are ultimately tied to subducted components. Comparisons with early Miocene alkali basalts provide evidence that subduction-related components first appeared in the mantle source under the Sierra de Chachahuén after 20 Ma.
Evolution of the late Miocene Chachahuén volcanic complex This evidence comes from the trace-element characteristics of the early Miocene basalts, which are like those of intraplate magmas that lack a subduction component (La/Ta < 12; Ba/Ta < 180; Ba/La < 20; Th/Hf ~ 0.5–0.6; Figs. 7 and 8A; Kay and Copeland, this volume, chapter 9). Higher Ta/Hf (~0.5; Fig. 9A) and La/Sm ratios at a given Th content (Fig. 9B) in the alkali basalts relative to the Chachahuén magmatic rocks also support derivation of the early Miocene magmas from an enriched oceanic-island basalt (OIB)–like mantle source little affected by fluid mobile elements. Other evidence for addition of a subducted crustal component into the mantle source after 20 Ma comes from higher εNd values and lower 87Sr/ 86Sr ratios in the alkali basalt than in the Chachahuén volcanic rocks (Fig. 10A). At the least, the early Miocene basalts provide no evidence for melting and assimilation of arc-like components into the magma source of the early Miocene. Support for near purging of the subducted components from the mantle source beneath the Sierra de Chachahuén by the late Pliocene comes from intraplate-like La/Ta (<15) ratios in alkaline basalts younger than 3 Ma (Fig. 8; Kay et al., this volume, chapter 2). Lower 87Sr/ 86Sr and 206Pb/ 204Pb ratios and higher εNd in these younger basalts compared to those in the Late Chachahuén group volcanic rocks (Fig. 10) also fit with the loss of a subducted crustal component in the mantle source. Intermediate ratios for Ba/La, Ta/Hf, and Th/Hf in the younger than 3 Ma basalts relative to the early Miocene and Early and Late Chachahuén group mafic rocks (Figs. 7B and 8A) signify either the residual effects of a subducted component in the lithosphere or an asthenospheric mantle source with these characteristics. The latter fits better with the relatively low εNd values in the basalts younger than 3 Ma (Fig. 10A; see Kay et al., this volume, chapter 2). Chemical Similarities of the Chachahuén Magmas to Chilean Flat-Slab Backarc Magmas Further support for eruption of the Chachahuén volcanic complex over a shallowly dipping subducting slab comes from chemical similarities with volcanic rocks erupted in the ca.7.9–4.5 Ma Pocho field near 31.5°S (Kay and Gordillo, 1994) and the ca. 6–2 Ma San Luis field hear 33.5°S (see Kay and Mpodozis, 2002) in the Chilean flat-slab region (28°S–33°S latitude; Fig. 1). The eruption of hornblendebearing basaltic to dacitic magmas up to 750 km east of the Chile Trench over the Chilean flat-slab has been associated with a rapid stage of shallowing of the Nazca plate starting at ca. 8 Ma (e.g., Kay and Abbruzzi, 1996). Chemical similarities between the Pocho and the Early and Late Chachahuén group volcanic rocks include: (1) plotting in the high-K to shoshonitic fields on a K2O-SiO2 diagram (Fig. 5), (2) REE patterns with overlapping La/Yb, La/Sm, and Sm/Yb ratios (Fig. 7), and (3) ratios of La/Ta > 25, Ba/La > 20, and Ta/Hf < 0.3 (Figs. 8 and 9). The Pocho volcanic rocks are also like the Chachahuén volcanic complex rocks in showing an older-to-younger temporal
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trend toward a more arc-like character as La/Ta, Ba/Ta, and Th/Hf ratios increase from the high-K to the shoshonitic series (Kay and Gordillo, 1994). Isotopic ratios in the Pocho rocks are different from those in the Chachahuén complex, but trends are similar in that 87Sr/ 86Sr ratios increase and εNd values decrease as La/Ta ratios increase (Figs. 8 and 10A). Lower εNd values and 206Pb/ 204Pb ratios in the Pocho volcanic rocks show that the Pocho magmas were derived from an isotopically distinct source (Figs. 10A and 10B). Timing and Style of Uplift of the Sierra de Chachahuén Evidence supporting the uplift of the Sierra de Chachahuén after the early Miocene and before the Pliocene comes from the eruption of the early Miocene alkali basalt flows in at least a mildly extensional setting (see Pérez and Condat, 1996; Kay, 2002; Kay et al., 2005) and the fact that alkali basalts younger than 3 Ma flowed down the broad shoulders of the range. Support for a predominantly late Miocene uplift comes from the distribution of the late Miocene volcanic rocks and the maps and discussion in Pérez and Condat (1996) that show that faults and fractures affecting them (Fig. 3) are related to both caldera formation and regional compression. The presence of amphibole phenocrysts and the crystallization sequence in the Early and Late Chachahuén group volcanic rocks support a period of magma evolution at lower- to mid-crustal depths consistent with crustal storage of magmas favored by a contractional stress regime. Overall, the topographic relief of the Sierr1a de Chachahuén is best explained by a combination of uplift on high-angle contractional faults and magmatic addition. Other support for late Miocene uplift of the Sierra de Chachahuén comes from evidence for contractional deformation in the backarc (Fig. 2) between 12 and 3 Ma (see Kozlowski et al., 1996; Kay, 2002; Cobbold and Rossello, 2003; Kay et al., this volume, chapter 2). Starting in the west, evidence for renewed uplift of the Cordillera del Viento and deformation of the middle Miocene Trapa Trapa arc volcanic sequences in the late Miocene comes from fission-track cooling ages (Burns, 2002). Evidence for latest middle to late Miocene compressional deformation near 35°S is documented by ages and field relations presented by Baldauf (1997). Moving eastward, a crustal cross section from the Cordillera del Viento to the Sierra de Huantraico in Zapata et al. (1999) is consistent with late Miocene compressional inversion of normal faults uplifting the Tromen block prior to Pliocene extension. A time window for uplift of the Tromen block comes from geochemical evidence that fits with crustal thickening related to crustal shortening between the emplacement of the Cerro Negro andesite at ca. 12 Ma and the eruption of Pliocene Tromen region lavas (Kay et al., this volume, chapter 2). Other structural evidence for uplift of the Tromen massif at this time was discussed by Cobbold and Rossello (2003). Similarities between the Chachahuén volcanic complex in the Sierra de Chachahuén and the late Miocene Pocho volcanic
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field in the Sierra de Cordoba in the Pampean Ranges over the Chilean flat-slab are consistent with a relation between magmatism and uplift in both regions. Importantly, eruptions of the Pocho lavas can be shown to overlap the initial stages of uplift of the Pampean Sierra de Cordoba range on high-angle reverse faults (see Kay and Gordillo, 1994; Ramos et al., 2002). Both magmatism and uplift have been correlated with a rapid stage of shallowing of the subducting Nazca plate under the modern Chilean flat-slab region (see Kay et al., 1991, 1999; Kay and Abbruzzi, 1996). Likewise, eruption of the Chachahuén volcanic complex and uplift of the Sierra de Chachahuén are consistent with an episode of shallow subduction under the Neuquén Basin as shown in Figure 11 and discussed in the next section. A Shallow Subduction Model for the Setting of the Chachahuén Volcanic Complex A short-lived shallow subduction zone under the Neuquén Basin provides a rationale for explaining the transient appearance of a basaltic to rhyodacitic volcanic complex with a subduction-
Figure 11. (A) Cartoon showing the lithospheric-scale cross section at ca. 8–5 Ma across the Chilean flat-slab near 30°S that Kay and Abbruzzi (1996) used to explain the distribution of late Miocene volcanic rocks and deformation in that region. (B) Cross section across the Neuquén Basin at ca. 8–5 Ma using a shallowly dipping slab to explain the distribution of late Miocene volcanic rocks and deformation in a profile near 36°S. Similarities and differences between the two models are discussed in the text. Locations of the cross sections are shown in Figure 1.
zone signature in the Sierra de Chachahuén, uplift of the Sierra de Chachahuén, and Miocene compressional deformation across the backarc. A lithospheric-scale cross section at ca. 8–5 Ma in Figure 11 shows a shallowly subducting slab under the region just as it reached its maximum development. Such a model can explain inversion of older extensional structures causing uplift of structural blocks across the region. Uplift is in accord with fluids released from the subducting slab hydrating and weakening the continental lithosphere prior to compressional deformation, as has been argued in other Andean shallow subduction zones (James and Sacks, 1999; Kay et al., 1999; Ramos et al., 2002). The transient nature of the shallowly subducting slab is based on changes in structural style and the absence of arc-like chemistry in both early Miocene and post–early Pliocene alkaline lavas. Subsequent steepening of the slab is inferred to be responsible for widespread post–late Pliocene basaltic magmatism, which is interpreted to result from mantle melting as asthenospheric mantle filled the region of the former shallow subduction zone (Kay et al., this volume, chapter 2). Comparisons with the latest Miocene lithospheric-scale cross section across the Chilean flat-slab near 30°S from Kay et al. (1991) and Kay and Abbruzzi (1996) are instructive. These authors argued that the latest Miocene shape of the subducting Nazca plate in the Chilean flat-slab region was fairly similar to the modern shape, which is inferred from the distribution of earthquake epicenters (Cahill and Isacks, 1992; Pardo et al., 2002). The shallow part of the late Miocene slab under the Neuquén Basin is shown as less pronounced, since the Chachahuén volcanic complex is ~500 km from the trench, whereas the Pocho volcanic field over the Chilean flat-slab is more than 700 km from the trench. A further difference is that post–early Miocene backarc shortening across the region is on the order of 37 km near 36°S (Zapata et al., 1999) as opposed to >150 km over the Chilean flat-slab (Allmendinger et al., 1990; Ramos et al., 2002). Crustal thickening in compensation for this crustal shortening has been used to explain why crustal thicknesses in the Main Cordillera are on the order of 60–70 km near 30°S and closer to 40–45 km near 36°S (e.g., Introcaso et al., 1992). In analogy with the chemically similar Pocho volcanic field over the Chilean flat-slab, the Chachahuén volcanic complex is inferred to have erupted ~180 km above the slab. As in the Chilean flat-slab region, backarc volcanic activity was spatially discontinuous. The nearest Miocene arc-like center was the ca. 12 Ma Cerro Negro center to the east (Fig. 11; Kay et al., this volume, chapter 2). The gap, like that between the Cerro Blanco and Pocho volcanics over the Chilean flat-slab, fits with discontinuous dehydration of the slab as different hydrous phases breakdown with depth (e.g., Poli and Schmidt, 1995). The K- and Ba-rich Chachahuén volcanic complex rocks, like those in the Pocho volcanic field, fit with a role for the breakdown of phlogopite, as is commonly argued for K-rich backarc
Evolution of the late Miocene Chachahuén volcanic complex magmas erupted far above the slab (see Kay and Gordillo, 1994). Another factor in explaining the high large ion lithophile element concentrations is relatively small partial melting percentages of the backarc mantle. CONCLUSIONS AND COMPLICATIONS The evolution of the ca 7.6 Ma to 4.8 Ma Chachahuén volcanic complex from an orthopyroxene-bearing silicic andesitic to rhyodacitic center with an intraplate to arc-like chemistry to a basaltic andesitic to dacitic hornblende-bearing nested caldera complex with an arc-like chemistry can be explained by a period of transient shallow subduction of the Nazca plate under the Neuquén Basin. The changing chemical characteristics of the lavas fit with a progressive increase in the influence of subducted components on mantle-derived magmas that intruded into the crust and increasingly affected the composition of that crust and the style of volcanism. A transient period of shallow subduction is supported by the intraplate, rather than arc-like character of both early Miocene and Pliocene basalts in the Sierra de Chachahuén along with the Miocene to Holocene magmatic and structural evolution of the Neuquén Basin (see Kay et al., this volume, chapter 2). The absence of backarc volcanic complexes with arc-like characteristics south of the Sierra de Chachahuén in the Neuquén Basin is consistent with the Cortaderas lineament (Fig. 2) marking the southern limit of the influence of shallow subduction (Kay et al., this volume, chapter 2). An important question is the cause of transient shallow subduction under the Neuquén Basin. A complicated issue is the relation of this shallowing to that under the Chilean flat-slab between 28°S and 33°S. The similarity in timing and chemistry of the Chachahuén and Pocho magmatic events and the associated uplift of block-faulted ranges fit with linked shallowing in both regions. In the Pliocene, the tectonic history of the two areas diverged, with slab flattening continuing in the Chilean flat-slab, but reverting to a steeper subduction zone under the Neuquén Basin. Shallowing in both regions cannot be explained by the popular model that links the shallowing of the Chilean flat-slab with subduction of the Juan Fernandez Ridge on the Nazca plate near 33°S (Fig. 1; e.g., Yáñez et al., 2001; Gutscher et al., 2000). One possibility is that the common timing is a coincidence, and a smaller segment of thickened oceanic crust subducted beneath the Neuquén Basin at the same time. However, the story could be more complex as the intervening region between 33°S and 36°S is in a backarc position to where the magmatic front has migrated eastward in the last 20 m.y., with peaks at ca. 20–16 Ma and ca. 8–5 Ma (e.g., Kay et al., 2005; Fig. 1). This arc migration, along with the formation of the Cerro Nevado volcanic complexes near 35°S (Fig. 1; Bermúdez, 1991), is consistent with more moderate Miocene to early Pliocene transient shallow subduction in the region that links the Chilean flat-slab with the Neuquén Basin.
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In a larger framework, the inferred period of shallower subduction under the Neuquén Basin happened at a time of major tectonic changes along much of the Andean margin (e.g., Kay et al., 1999). On a final note, Tebbens and Cande (1997) noted that changes in the evolution of the southeastern Pacific Ocean at Chron 3A at ca. 5.5 Ma are synchronous with an extended interval of plate boundary reorganization in the northeast Pacific. ACKNOWLEDGMENTS Funding for this study was provided by a grant from Repsol YPF petroleum company, with additional support from the U.S. National Science Foundation (NSF) grant 00–87515. Special thanks are extended to Daniel Ragona for his role in the field work, help with sample preparation, assistance in processing satellite data, and insights into structural problems in the Sierra de Chachahuén. Thanks are also extended to Facundo Fuentes for processing satellite images and to Linda Godfrey for help with isotopic analyses. Discussions with Tomas Zapata, Steven Sparks, Víctor Ramos, and particularly Robert Kay were beneficial in the development of the ideas presented here. Todd Feeley and Beatriz Coira are thanked for helpful reviews that improved the quality of the paper. APPENDIX 1. CHACHAHUÉN VOLCANIC COMPLEX SAMPLE LOCATIONS AND DESCRIPTIONS Cerro de Las Vizcachas DR44. Rhyodacite ignimbrite from Cerro de Las Vizcachas, near Puesto Castillo, 37°05′45.4″S, 69°00′14.3″W. Petrography: Silicic ignimbrite with biotite. DRC23. Dacite porphyry, altered, below small waterfall in section in sketch of diorite in Holmberg (1962), 37°05.10′S, 69°00.1′W. DRC24. (a) pink rhyodacite ignimbrite, (b) hornblende-bearing andesite, 37°05.10′S, 69°00.1′W. Sierra de Chachahuén Region Cañadón del Camino South of Puesto Zuñiga, 37°6.08′S, 68°52.55′W, 1300 ± 34 m DR4. Basalt flow with prominent euhedral clinopyroxene phenocrysts up to several centimeters across. DR4N. Clinopyroxene-bearing basalt flow along Quebrada del Camino road, near DR4. DR4B. Clinopyroxene-bearing basalt flow, north along Quebrada del Camino road, close to DR4N. Cerro “Condor,” 37°4.36′S, 68°53.18′W, 1450 ± 30 m at Base DR5. Mafic hornblende andesite from massive, columnar jointed flow near top of Cerro “Condor,” unit 5 south of
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Aguada del Tunduco. Petrography: Phenocrysts of clinopyroxene, plagioclase, and amphibole with oxidized rims in devitrified groundmass. DR6. Andesite clast on north side of Cerro “Condor,” crest of small hill, block and ash deposit, unit 4. DR7. Andesite clast near base of Cerro “Condor,” block and ash in narrow quebrada, unit 4. Cerro Chachahuén DR8. Dacite “ intrusive” on northwest side of Cerro Chachahuén, east of Cretaceous Neuquén Group south of Puesto La Cortaderita, 37°4.28′S, 68°51.45′W, ~1450 m. Petrography: Clinopyroxene, plagioclase, amphibole phenocrysts in devitrified groundmass; feldspars have melt in cores. DR9. Hornblende andesite clast in agglomerate above DR8, unit 3, northwest side of Cerro Chachahuén, 37°4.33′S, 68°51.45′, 1726 ± 5 m. Petrography: Phenocrysts of hornblende with oxidized rims, clinopyroxene, plagioclase, with melt textures. DR10. Clinopyroxene-bearing basaltic andesite flow, about ~6 m thick, top of Cerro Chachahuén, Cerro Chachahuén volcanic unit, 37°04.55′S, 68°51.00′W, 1929 ± 20 m. DR11. Rose-colored pyroxene-bearing mafic andesite flow below and west of top of Cerro Chachahuén, Cerro Chachahuén volcanic unit, 37°04.55′S, 68°51.18′W, 1901 ± 37 m. Petrography: Clinopyroxene, olivine, and plagioclase phenocrysts; clots of clinopyroxene and plagioclase, calcite in vesicles. DR12. Pink pyroxene-bearing vesicular mafic andesite flow in middle part of northwest side of Cerro Chachahuén, Cerro Chachahuén volcanic unit, 37°04.42′S, 68°51.42′W, 1648 ± 39 m. Petrography: Oxyhornblende phenocrysts with oxidized rims, feldspar with melt textures, sparse clinopyroxene phenocrysts in devitrified groundmass. DRC13. Basaltic andesite, long flow from Cerro Chachahuén east of Quebrada del Camino, 37°06.6′S, 68°51.24′W. Cerros Bayo and Boina Region DR2. Dacitic vitrophyre from agglomerate, unit 2, just north of Cerro Bayo, 36°59.52′S, 68°53.29′W, 1450 ± 30 m. DR13. Hornblende-bearing mafic andesite flow, northeast edge of quebrada southeast of Cerro Bayo, some alteration, unit 4, 37°03.36′S, 68°51.06′W, 1403 ± 20 m. DR14. Main hornblende-bearing andesite unit composing Cerro Bayo, moderate alteration, from above quebrada on west side, 37°02.98′S, 68°51.57′W, 1434 ± 37 m. Petrography:
Euhedral green amphibole and plagioclase phenocrysts, complex crystal aggregates, magmatic titanite, fine-grained groundmass. DR15. Dacitic clast from agglomerate, east side of arroyo west of Cerro Bayo, unit 3 or 4, 37°2.94′S, 68°52.01′W. DR16. Andesitic dike with prominent chill margin, columnar jointed, cuts DR15, 37°2.94′S, 68°52.01′W. DR20. Dacitic vitrophyre in unit 2 above early Miocene basalt, northeast of Puesto Cortaderita, 37°3.72′S, 68°51.74′W, 1396 ± 40 m. DR21. Mafic andesitic intrusive, cuts unit 3/4 agglomerate in quebrada, 37°3.30′S, 68°51.75′W, 1480 ± 37 m. Petrography: Hornblende phenocrysts with oxidized rims, plagioclase with melt textures in groundmass full of small feldspar laths. DR22. Basaltic clast in reddish agglomerate near top of Cerro Boina, unit 5, 37°3.26′S, 68°52.35′W, 1881 ± 88 m. DR23. Dacitic pumice in ignimbrite, south side of Cerro Boina, either unit 2 or 3, 37°03.3′S, 68°52.3′W, 1580 m. DR24a. Basaltic andesite clast in debris flow, clasts up to a meter across, south side Cerro Boina, unit 3/4, near 37°3.26′S, 68°52.35′W. Petrography: Phenocrysts of zoned green clinopyroxene, large amphibole with clinopyroxene cores, and plagioclase with melt in cores in fine-grained groundmass, calcite in vesicles. Cerro Corrales Region DR24b. Mafic andesite clast in unit 6 above contact with unit 5, halfway up south side of Cerro Corrales above Puesto Zuñiga. DR25. Andesitic pumice from pumice-rich ignimbrite, unit 5 just below unit 6 on south side of Cerro Corrales, pumice up to 25 cm across. DR26. Mafic andesite clast in agglomerate above DR25, unit 6 south of top of Cerro Corrales, 37°3.94′S, 68°55.04′W, 1694 ± 80 m. DR27a and b. Hornblende andesite dike with large gabbroic xenoliths; dike essentially forms a vertical wall; samples are from finer-grained margin of dike below and to south west of DR26, 37°04.1′S, 68°55.4′W, 1710 m. Petrography: Very large hornblende crystals surrounding clinopyroxene and titanomagnetite, fine-grained groundmass. DR28. Pinkish andesite with acicular hornblende crystals, south side of Cerro Corrales, 37°4.21′S, 68°55.31′W, 1698 ± 45 m. DR29. Amphibole-bearing basaltic andesite dike cutting Cretaceous Agrio Formation southeast of Cerro Corrales, 37°4.69′S, 68°54.62′W, 1403 ± 35 m. Petrography: Hornblende, plagioclase and minor clinopyroxene phenocrysts in fine-grained devitrified groundmass.
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Cerro Campanarío and Cerro de la Cabeza Region DR30. Andesitic Complejo Volcanico, collapsed dome in middle of caldera, southeast of Cerro Campanarío, 37°2.45′S, 68°53.85′W, 1711 ± 39 m. Petrography: Feldspar, orthopyroxene, and oxidized amphibole phenocrysts in devitrified glassy matrix. DR31. Mafic andesitic dike, N-S–trending in unit 3, ~3 m wide, east of quebrada west of high point of Cerro de la Cabeza, 37°01.64′S, 68°53.70′W, 1613 ± 49 m. DR33. Mafic andesite from “intrusive” in unit 3 northeast of basalt cone, 37°00.69′S, 68°53.19′W, 1499 ± 28 m. DR34. Layered, dipping rhyodacite layer, Cerro de la Gloria intrusive, south side along quebrada (Rio del Agua Lopina), 36°59.41′S, 68°52.11′W, 1370 ± 62 m. DR35. Dacite similar to DR34, near altered hot spring deposits, layered, west of Cerro de la Gloria, 36°59.21′S, 68°52.55′W, 1469 ± 42 m. DR36. Dacitic block from block and ash deposits, Complejo Volcanico unit, Puesto Isaac, 37°01.74′S, 68°54.52′W, 1429 ± 59 m. DR37. Mafic andesite dike, ~7 m wide, cuts block and ash deposits, <100 m southeast of DR36, 37°01.74′S, 68°54.52′W, 1429 ± 59 m.
Locality of Ch19 dated at 5.95 ± 0.3 Ma by Pérez and Condat (1996), 37°03.3′S, 68°52.3′W, ~1580 m. DR52. Basalt flow near Cerro Bayo in western Sierra de Chachahuén, south of Cerro de la Montura, 37°02.53′S, 68°58.9′W. Petrography: Oxidized amphibole, plagioclase and clinopyroxene phenocrysts, clinopyroxene aggregates, ophitic texture. DR53. Dacitic intrusive west of Cerro Bayo in western Sierra de Chachahuén, south of Cerro de la Montura, 37°01′57.4″S, 68°58′ 46.6″W. DR54. Dacitic ignimbrite, Cerro de los Rodados, 37°00.44′S, 68°56.84′W. DR55. Rhyodacitic vitrophyre in unit 2, northeast of Cerro Bayo, 37°02.04′S, 68°50.78′W.
Western Sierra de Chachahuén DR39. Basaltic andesite flow, unit 6, Cerro de la Montura, south of road in Río Café, 37°00.68′S, 68°57.67′W, 1666 ± 48 m. DR40. Basaltic andesite flow above DR39, unit 6, uppermost flow from Cerro de la Montura, south of road in Río Café, 37°00.75′S 68°57.87′W, 1534 ± 31 m. Petrography: Rims of oxidized amphibole on clinopyroxene phenocrysts, plagioclase phenocrysts. DR45. Mafic andesite flow, unit 6, north of Cerro Corrales, 37°03′49.7″S, 68°54′59.7″W. DR46. Mafic andesite clast in agglomerate, unit 5, north of Cerro Corrales, 37°03′40.7″S, 68°55′13.7″W. DR47. Mafic andesite flow, northeast of Cerro Corrales, 37°02′47.1″S, 69°55′08.1″W. Petrography: Amphibole with oxidized rims, clinopyroxene and plagioclase phenocrysts in devitrified fine-grained matrix. DR48. Silicic andesite from “Great Dike” of Pérez and Condat (1996), south of Cerro Ureta. Petrography: Green amphibole and large plagioclase phenocrysts in devitrified groundmass, 37°03.5′S, 68°55.1′W, ~1950 m. DR49. Mafic andesite flow, south of Puesto Zuñiga, isolated unit 5 outcrop west of Cañadón del Camino, west of DR4.
Chachahuén Volcanic Complex—Cerro Chachahuén DRC7. Unit 2 ignimbrite complex along road just east of Puesto Zuñiga, (a) reddish dacite clast from upper agglomerate, (b) dacitic pumice from lower unit, (c) gray dacitic clast in pyroclastic flow above pumice layer, 37°04.34′S, 68°53.87′W, 1464 ± 34 m. DRC8. Dacitic vitrophyre, unit 2 north side of Cerro Bayo, 37°02.36′S, 68°52.11′W, 1607 m. DRC9a. Mafic andesite dike west of Cerro de la Cabeza in quebrada, 37°01.2′S, 68°53.5′W, ~1580 m. DRC9. Basalt flow east of Quebrada del Camino, short distance east of DR4, near road to Puesto Sosa, 37°06.1′S, 68°52.356′W, ~1320 m.
Arroyo los Golpeados Region near 37°05.44′S, 68°54.72′W DRC3. Dacitic ignimbrite, unit 2 on east side of Arroyo los Golpeados. DRC5. Andesitic (a) clast and (b) pumice in unit 2 just above early Miocene basalt. DRC6. Rhyodacitic welded tuff in Arroyo los Golpeados. “Fluidal andesite” Ch5 dated at 6.4 ± 0.3 Ma by Pérez and Condat (1996),
Base of South Side of Cerro Corrales DRC10. (a) Dacitic pumice from steep white cliff, (b) basaltic clast in debris flow nearby, 37°04.35′S, 68°54.56′W. DRC11. Mafic andesitic clast in agglomerate, 37°04.35′S, 68°54.24′W. DRC12. Hornblende-bearing dike in quebrada near base of Cerro Corrales north of Puesto Zuñiga, east of DRC10A, above early Miocene basalt, 37°04.35′S, 68°54.24′W.
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APPENDIX 2: 40Ar/ 39Ar AGE SPECTRUM
Figure A1. Age spectra for samples in Table 1 dated by the 40Ar/ 39Ar step-heating method (Mitchell, 1968) at the University of Houston in the laboratory of Peter Copeland.
APPENDIX 3. ANALYTICAL METHODS Major-element chemical analyses were performed on fused glasses using the JEOL 733 Superprobe in the Cornell University Materials Science Center (CCMR). Samples with <60% SiO2 were melted in a molybdenum strip furnace in an Ar atmosphere, whereas samples with >60% SiO2 were mixed with a meta-borate flux and melted in carbon crucibles in air at 1000 °C for 30 min in a muffle furnace. Microprobe analyses were preformed at 15 kV with a current of 15 A using wavelengthdispersive spectrometers. Reported analyses are averages of five 20–30-μm-diameter spots. Smithsonian standard Juan de Fuca glass was used as a secondary standard. Analyses on fluxed glasses were normalized to 100%. Trace-element analyses were preformed by instrumental neutron activation analyses (INAA) at the Ward Reactor Center at Cornell University. Powdered samples (~0.5 g) were sealed in high-purity silica glass
tubes and irradiated in the Cornell Triga reactor at a neutron flux of 5 × 1013 neutrons cm–2 s–1 for 2 h. Gamma-ray spectra were acquired using an ORTEC intrinsic Ge coaxial detector (20% efficiency; resolution of 0.7 KeV at 1.33 MeV), a Nuclear Data ADC, and a multichannel analyzer. Samples were counted for a minimum of 2 h (up to 10 h), 7 and 40 d after irradiation. Data were reduced using the Cornell data reduction program. Corrections for peak interferences were made on Ce (for Fe), Nd (for Br), Tb (for Th), Eu (for Ba), Lu (for U), and Yb (for Th). Whole-rock FeO concentrations were used as internal flux monitors, and trace-element concentrations are proportional to FeO concentrations. FeO concentrations based on INAA counting data and Na2O concentrations determined by INAA data were cross-checked against whole-rock analyses to ensure analyses were accurately matched. Priority was given to Na2O concentrations from INAA analyses. Reported analyses are volatile free because they were proportioned to Fe concentra-
Evolution of the late Miocene Chachahuén volcanic complex tions from volatile-free analyses on glasses. All isotopic analyses were done at Cornell University on a VG Sector thermal ionization mass spectrometer (TIMS). Sr isotopic ratios were measured on W single filaments using a quadruple-collector dynamic procedure. Ratios were normalized to an 87Sr/ 86Sr ratio of 0.1194. Nd isotopic analyses were measured on single Re filaments using a quintuple-collector dynamic procedure. Average analytical values for 87Sr/ 86Sr for NBS987 were 0.710221 (±0.000044), and 143Nd/ 144Nd ratios for La Jolla are 0.511888 (±0.000055). The 143Nd/ 144Nd ratios were corrected to a value of 0.5118624 for the La Jolla standard, and εNd was calculated based on a value of –15.15 for La Jolla. Pb isotopic ratios were corrected for mass fractionation based on ratios of 206Pb/ 204Pb = 16.931, 207Pb/ 204Pb = 15.485, and 208Pb/ 204Pb = 36.681 measured on Pb standard NBS SRM981. REFERENCES CITED Allmendinger, R.W., Figueroa, D., Snyder, D., Beer, J., Mpodozis, C., and Isacks, B.L., 1990, Foreland shortening and crustal balancing in the Andes at 30°S latitude: Tectonics, v. 9, p. 789–809. Baldauf, P.E., 1997, Timing of the uplift of the Cordillera Principal, Mendoza Province, Argentina, [Ph.D. thesis]: Washington, D.C., George Washington University, 356 p. Bermúdez, A., 1991, Sierra del Nevada, el limite oriental del arco volcánico Neógeno entre los 35°30′–36°L.S., Argentina, in VI Congreso Geológico Chileno: Actas, v. 1, p. 318–322. Bermúdez, A., and Delpino, D., 1989, La provincia basáltica Andesina cuyana (35°–37°S): Asociación Geológica Argentina Revista, v. 44, p. 28–34. Bermúdez, A., Delpino, D., Frey, F., and Saal, A., 1993, Los basaltos de retroarco extraandinos. in Ramos, V.A., ed., Geología y recursos naturales de Mendoza, in XII Congreso Geológico Argentino, Mendoza: Relatorio, p. 173–195. Burns, W.M., 2002, Tectonic and depositional evolution of the Tertiary Cura Mallín basin in the southern Andes (36.5 to 38°S lat.) [Ph.D. thesis]: Ithaca, New York, Cornell University, 218 p. Cahill, T.A., and Isacks, B.L., 1992, Seismicity and shape of the subducted Nazca plate: Journal of Geophysical Research, v. 97, p. 17,503–17,529. Cobbold, P.R., and Rossello, 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. González Díaz, E.F., 1979, Descripción geológica de la hoja 31d, La Matancilla, Argentina: Buenos Aires, Dirección Nacional de Geología y Minería, Boletín, escala 1:200,000, 96 p. Gutscher, M., Maury, R., Eissen, J.-P., and Bourdon, E., 2000, Can slab melting be caused by flat subduction?: Geology, v. 28, p. 535–538, doi: 10.1130/ 0091-7613(2000)028<0535:CSMBCB>2.3.CO;2. Hickey, R.L., Frey, F.A., Gerlach, D.C., and Lopez-Escobar, L., 1986, Multiple sources for basaltic arc rocks from the Southern Volcanic Zone of the Andes (34–41°S): Trace element and isotopic evidence for contributions from subducted oceanic crust, mantle and continental crust: Journal of Geophysical Research, v. 91, p. 5963–5983. Holmberg, E., 1962, Descripción geológica de la hoja 32-d, Chachahuén, Provincias de Neuquén y Mendoza: Buenos Aires, Argentina, Dirección Nacional de Geología y Minería, Boletín, no. 91, 70 p. Introcaso, A., Pacino, M.C., and Fraga, H., 1992, Gravity, isostasy, and Andean crustal shortening between 30 and 35°S: Tectonophysics, v. 205, p. 31–48, doi: 10.1016/0040-1951(92)90416-4. James, D.E., and Sacks, S.I., 1999, Cenozoic formation of the Central Andes: A geophysical perspective, in Skinner, B., ed., Geology and ore deposits
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of the Central Andes: Society of Economic Geology Special Publication 7, p. 1–25. Kay, S. M., 2001, Geochemical evidence for a late Miocene shallow subduction zone in the Andean Southern Volcanic Zone near 37° S latitude: Eos (Transactions, American Geophysical Union) v. 81, abstract V12C-099. Kay, S.M., 2002, Tertiary to Recent transient shallow subduction zones in the Central and Southern Andes, in XV Congreso Geológico Argentina: Actas, v. 6, contribution 237, CD-ROM, ISBN 987–20190–1-0. Kay, S.M., and Abbruzzi, J.M., 1996, Magmatic evidence for Neogene lithospheric evolution of the central Andean “flat-slab” between 30° and 32°S: Tectonophysics, v. 259, p. 15–28, doi: 10.1016/0040-1951(96)00032-7. Kay, S.M., and Copeland, P., 2006, this volume, 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, doi: 10.1130/2006.2407(09). Kay, S.M., and Gordillo, C.E., 1994, Pocho volcanic rocks and the melting of depleted continental lithosphere above a shallowly dipping subduction zone in the Central Andes: Contributions to Mineralogy and Petrology, v. 117, p. 25–44, doi: 10.1007/BF00307727. Kay, S.M., and Mancilla, O., 2001, Neogene shallow subduction segments in the Chilean/Argentine Andes and Andean-type margins: Geological Society of America Abstracts with Programs, v. 34, no. 6, p. A-156. Kay, S.M., and Mpodozis, C., 2002, Magmatism as a probe to the Neogene shallowing of the Nazca plate beneath the modern Chilean flatslab: Journal of South American Earth Sciences, v. 15, p. 39–58, doi: 10.1016/S08959811(02)00005-6. 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., Mpodozis, C., Ramos, V.A., and Munizaga, F., 1991, Magma source variations for mid-late Tertiary magmatic rocks associated with a shallowing subduction zone and a thickening crust in the central Andes (28 to 33°S), in Harmon, R.S., and Rapela, C.W., eds., Andean magmatism and its tectonic setting: Geological Society of America Special Paper 265, p. 113–137. Kay, S.M., Mpodozis, C., and Coira, B., 1999, Magmatism, tectonism, and mineral deposits of the Central Andes (22°–33°S latitude, in Skinner, B., ed., Geology and ore deposits of the Central Andes: Society of Economic Geology Special Publication 7, p. 27–59. Kay, S.M., Godoy, E., and Kurtz, A., 2005, Episodic arc migration, crustal thickening, subduction erosion, and Miocene to Recent magmatism along the Andean Southern Volcanic Zone margin: Geological Society of America Bulletin, v. 117, p. 67–88, doi: 10.1130/B25431.1. Kay, S.M., Burns, W.M., Copeland, P., and Mancilla, O., 2006, this volume, 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, doi: 10.1130/2006.2407(02). Kozlowski, E.E., Cruz, C.E., and Sylvan, C.A., 1996, Geología estructural de la zona de Chos Malal, Cuenca Neuquina, Argentina, in XIII Congreso Geológico Argentino y III Congreso de Exploración de Hidrocarburos: Actas, v. I, p. 15–26. Linares, E., and González, R., 1990, Catalogo de edades radiométricas de la Republica Argentina 1957–1987: Buenos Aires, Publicaciones Especiales de la Asociación Geológica Argentina, serie B, no. 19, 628 p. Lipman, P.W., Dungan, M.A., Brown, L.L., and Deino, A., 1996, Recurrent eruption and subsidence of the Platoro caldera complex, southeastern San Juan volcanic field, Colorado: New tales from old tuffs: Geological Society of America Bulletin, v. 108, p. 1039–1055, doi: 10.1130/00167606(1996)108<1039:REASAT>2.3.CO;2.
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Luhr, J.F., Carmichael, I.S.E., and Varekamp, J.C., 1984, The 1982 eruptions of El Chichon Volcano, Chiapas, Mexico: Mineralogy and petrology of the anhydrite-bearing pumices: Journal of Volcanology and Geothermal Research, v. 23, p. 69–108, doi: 10.1016/0377-0273(84)90057-X. Mitchell, J.G., 1968, The argon-40/argon-39 method for potassium-argon age determination: Geochimica et Cosmochimica Acta, v. 32, p. 781–790. Miyashiro, A., 1974, Volcanic rock series in island arcs and continental margins: American Journal of Science, v. 274, p. 321–355. Moore, G., and Carmichael, I.S.E., 1998, The hydrous phase equilibria (to 3 k-bars) of an andesite and basaltic andesite from western Mexico: Constraints on water content and conditions of phenocryst growth: Contributions to Mineralogy and Petrology, v. 130, p. 304–319, doi: 10.1007/ s004100050367. Muñoz, J., Stern, C.R., Bermúdez, A., Delpino, D., Dobbs, M.F., and Frey, F.A., 1989, El volcanismo Plio-Cuaternario a través de los 34°–39°S de los Andes: Asociación Geológica Argentina Revista, v. 44, p. 270–286. Pankhurst, R.J., and Rapela, C.W., 1995, Production of Jurassic rhyolite by anatexis of the lower crust of Patagonia: Earth and Planetary Science Letters, v. 134, p. 23–36, doi: 10.1016/0012-821X(95)00103-J. 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. Pérez, M.A., and Condat, P., 1996, Geología de La Sierra de Chachahuén, Area CNQ-23, Puelen: Buenos Aires, Argentina, Geólogos Asociados, S.A., report to YPF, 82 p. Poli, S., and Schmidt, M.W., 1995, H2O transport and release in subduction zones: Experimental constraints on basaltic and andesitic systems: Journal of Geophysical Research, v. 100, p. 22,299–22,314, doi: 10.1029/ 95JB01570. Ramos, V.A., 1978, Estructura, in Rolleri, E.O., ed., Geología y Recursos Naturales de la Provincia del Neuquén. VII° Congreso Geológico Argentino (Neuquén), Relatorio, p. 99–118. Ramos, V.A., Cristallini, E., and Pérez, D.A., 2002, The Pampean flat-slab of the Central Andes: Journal of South American Earth Sciences, v. 15, p. 59–78, doi: 10.1016/S0895-9811(02)00006-8. Saal, A., 1994, Petrology and geochemistry of intra-back arc basalts from the Argentine Andes [Master’s thesis]: Boston, Massachusetts, Massachusetts Institute of Technology, 164 p.
Saal, A., Frey, F.A., Delpino, D., and Bermudez, A., 1993, Geochemical characteristics of alkalic basalts erupted behind the Andean volcanic front (35°–37°S); constraints on sources and processes involved in continental arc magmatism: Eos (Transactions, American Geophysical Union), v. 74, p. 43. Stern, C.R., 1991, Role of subduction erosion in the generation of Andean magmas: Geology, v. 19, p. 78–81, doi: 10.1130/0091-7613(1991)019 <0078:ROSEIT>2.3.CO;2. 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. Sun, S.-S., and McDonough, W.F., 1989, Chemical and isotopic systematics of oceanic basalts; implications for mantle composition and processes, in Saunders, A.D., and Norry, M.J., eds., Magmatism in the ocean basins: Geological Society [London] Special Publication 42, p. 313–345. Tebbens, S.F., and Cande, S.C., 1997, Southeast Pacific tectonic evolution from early Oligocene to Present: Journal of Geophysical Research, v. 102, p. 12,061–12,084, doi: 10.1029/96JB02582. 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., Suárez Soruco, R., and Welsink, H.J., eds., Petroleum basins of South America: Tulsa, Oklahoma, American Association of Petroleum Geologists Memoir 62, p. 383–402. Yáñez, G.A., Ramiro, C., von Huene, R., and Diaz, J., 2001, Magnetic anomaly interpretation across the southern central Andes (32°–34°S): The role of the Juan Fernández Ridge in the late Tertiary evolution of the margin: Journal of Geophysical Research, v. 106, p. 6325–6345, doi: 10.1029/ 2000JB900337. Zapata, T., Brissón, I., and Dzelalija, F., 1999, The structures of the Andean fold and thrust belt in relation to basement control in the Neuquén Basin: Boletín de Informaciones Petroleras, v. 16, p. 112–121. Zencich, S., 2000, Oil discovery in the volcan Auca Mahuida zone: Boletín de Informaciones Petroleras, v. 16, p. 18–29.
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Geological Society of America Special Papers Miocene to Quaternary deformation of the Guañacos fold-and-thrust belt in the Neuquén Andes between 37°S and 37°30 ′S Andrés Folguera, Víctor A. Ramos, Emilio F. González Díaz and Reginald Hermanns Geological Society of America Special Papers 2006;407;247-266 doi: 10.1130/2006.2407(11)
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Geological Society of America Special Paper 407 2006
Miocene to Quaternary deformation of the Guañacos fold-and-thrust belt in the Neuquén Andes between 37°S and 37 °30 ’S Andrés Folguera* Víctor A. Ramos* Laboratorio de Tectónica Andina, Universidad de Buenos Aires, Consejo Nacional de Investigaciones Científicas y Técnicas, Buenos Aires, Argentina Emilio F. González Díaz Universidad de Buenos Aires, Servicio Geológico Minero Argentino (SEGEMAR), Buenos Aires, Argentina Reginald Hermanns Geological Survey of Canada, Vancouver, British Columbia, Canada
ABSTRACT The fold-and-thrust belt of the Neuquén Andes between 37°S and 37°30’S can be divided into two sectors that show contrasting evolutionary styles. The eastern, or outer, section contains the relatively well-known Agrio and Chos Malal fold-and-thrust belts, which were active from the Late Cretaceous to the late Miocene. The western, or inner, sector contains the Guañacos fold-and-thrust belt, which is the subject of this paper. In contrast to the eastern-sector belts, contractional deformation in the Guañacos foldand-thrust belt began during the late Miocene and continues today. The Guañacos belt is particularly noteworthy for being out-of-sequence, in that the youngest deformation occurs in the western sector of the deformational belt in the Pleistocene volcanic arc, in contrast to the usual situation in the Andes, where the youngest brittle deformation occurs in the eastern sector. In detail, the Guañacos belt is located proximal to the present volcanic arc and coincides with the maximum heights of the Neuquén Andes. This thrust belt formed in response to tectonic inversion of an Oligocene-Miocene intraarc rift. Detailed structural traverses in three valleys and observations in two other valleys through this 60-km-wide thrust belt show that the youngest tectonic activity is concentrated in the easternmost 40 km. Over the last 5 m.y., folds and thrusts in the Guañacos belt have affected Pliocene to lower Pleistocene volcanic arc rocks as well as Quaternary deposits that are immediately east of the Upper Pleistocene to Holocene centers of the Southern volcanic zone. Among the volcanic rocks incorporated into the Guañacos fold-and-thrust belt are those of the Pliocene to Quaternary Trohunco caldera and the Los Cardos–Centinela volcanic center. The inversion of the Tertiary extensional structures in the Guañacos belt is considered to be mechanically linked with the La Laja strike-slip fault system in the intra-arc in Chile. Keywords: tectonic inversion, Tertiary extension, Andean Neotectonics, Late Miocene deformation. *E-mails:
[email protected];
[email protected]. Folguera, A., Ramos, V.A., González Díaz, E.F., and Hermanns, R., 2006, 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°S lat): Geological Society of America Special Paper 407, p. 247–266, doi: 10.1130/2006.2407(11). For permission to copy, contact
[email protected]. ©2006 Geological Society of America. All rights reserved.
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INTRODUCTION Several Tertiary to Quaternary fold-and-thrust belts with distinctive deformational styles have been identified in the Andes between 37°S and 40°S (Fig. 1). The region in which these belts occur constitutes the boundary between the southernmost Central Andes and the northernmost Patagonian Andes. This ~60-km-wide region is distinctive for having the lowest maximum elevations (2500 m) in the Andes between Colombia and Tierra del Fuego (Fig. 1). In detail, this region, which is known as the Neuquén Cordillera, is formed by two parallel fold-and-thrust belts. The eastern one consists of the relatively well-studied Agrio fold-and-thrust belt in the south and the Chos Malal fold-and-thrust belt in the north (e.g., Ramos, 1977; Kozlowski et al., 1996; Zapata et al., 2002; Mosquera and Ramos, this volume, chapter 5). The western belt, which is on the eastern slope of the Neuquén Andes, is here designated the Guañacos fold-and-thrust belt (Fig. 2). The focus of this paper is on the deformational history of the Guañacos fold-and-thrust belt. Emphasis is placed on the
post-Miocene deformational activity, since the Miocene activity has been described by Burns (2002) and Burns et al. (this volume, chapter 8). The nature of the young deformation in this belt is accessible through structural data gathered in five transects on the eastern slope of the Neuquén Andes (Fig. 2A). The observations show that a series of 10–12 emergent thrusts forms an eastward-propagating deformational front that affects the Pliocene to Quaternary calderas and stratovolcanoes of the main volcanic arc (Fig. 2) (Rovere, 1993, 1998). The easternmost thrusts have undergone Quaternary activity, which is considered to be associated with the landslides that produced the variable amounts of landscape denudation in this region (Fig. 3). TECTONIC SETTING AND REGIONAL GEOLOGY The western part of the Neuquén Andes between 36° and 38°S is covered by Upper Pleistocene to Holocene volcanic rocks erupted at the modern volcanic arc front. Contemporaneous retro-arc volcanic rocks to the east overlie late Pliocene to early Pleistocene arc volcanic rocks that have been interpreted as
Figure 1. Map shows the Guañacos, Agrio, and Chos Malal fold-and-thrust belts (FTB) relative to the active volcanoes (Vn) of the Southern volcanic zone arc and other major structural features in the transition zone between the southern central Andes and the northern Patagonian Andes. The large box shows the region in the map in Figure 2. The interior box shows the region studied in detail in this paper.
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Figure 2. Maps putting the Huaraco, Guañacos, and Lileo, Reñileuvú, and Ñireco structural traverses discussed in this paper in context with the major: (A) structural and volcanic features, and (B) stratigraphic units in the Neuquén Andes. Box shows region studied in detail. The times of deformation and ages shown on the map in A come from regional studies, fission-track analysis, and cooling ages (see Niemeyer and Muñoz, 1983; Burns, 2002; Kay 2002; Kay et al., this volume, chapter 2, and references within). The regional cross sections shown in Figure 14 are along the labeled line.
erupting in the eastern part of a broad frontal arc (Lara et al., 2001; Lara and Folguera, this volume, chapter 14). The arc front has not moved to the foreland during the Pliocene and Quaternary, making this segment distinct from much of the rest of the Andes, where the arc front has migrated eastward (Mpodozis and Ramos, 1989). The orogenic front in the Neuquén Andes is also distinctive from other parts of the Andes in having shifted westward into the modern retroarc area between the late Miocene and Quaternary (Ramos and Folguera, 2006; Zapata and Folguera, 2006). Evidence that the most recent deformational activity is occurring in the eastern (inner sector) of the retroarc is shown by the distribution of seismicity (Barrientos and Acevedo, 1992; Folguera et al., 2003; Zapata and Folguera, 2006) and
neotectonic deformation features (Folguera and Ramos, 2002; Folguera et al., 2004). Global positioning system (GPS) measurements (Kendrick et al., 1999) on the western boundary of the Central Depression in Chile (Fig. 2) further show that the forearc (hinterland) at 37°–38°S is being transported eastward to accommodate the contractional displacement in the retroarc (Fig. 1). The absence of contractional neotectonic activity in the Agrio and Chos Malal fold-and-thrust belts (Folguera et al., this volume, chapter 12) along with 40Ar/39Ar and fission-track data for Cretaceous to Miocene deformation in that region (Fig. 2; Burns, 2002; Kay, 2002; Kay et al., this volume, chapter 2) suggest that the overall displacement in the Andes at this latitude is absorbed relatively near the Quaternary arc front.
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Figure 3. Geological map between the Reñileuvú and Buraleo valleys on the eastern slope of the Andes showing the major structures and units in the easternmost portion of the Guañacos fold-and-thrust belt. Lines A to D indicate eastern sectors of the Huaraco, Rio Lileo, Guañacos, and Reñileuvú transects discussed in the text.
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Miocene to Quaternary deformation of the Guañacos fold-and-thrust belt The most prominent structural feature in the Holocene arc front at these latitudes is the La Laja dextral strike-slip system (Figs. 1 and 2) (Melnick et al., 2003; Melnick et al., this volume, chapter 4). The oldest age of deformation on this fault system, which shows evidence of Quaternary activity, is not well constrained. The symmetric geometry of Neogene structures on either side of the La Laja fault system suggests that it plays a role in the development of contractional deformational features on both the arc and the inner retroarc through a transpressional mechanics. Two morphostructural units occur along the eastern slope of the Neuquén Andes between 37° and 37°30′S (Figs. 1 and 2). The first, the Main Cordillera in the west, incorporates the Guañacos fold-and-thrust belt. The Upper Paleogene, Neogene, and Quaternary volcanic sequences in this region define a fairly broad and smooth topographic profile (Fig. 4). The second, to the east, incorporates the Agrio and Chos Malal fold-and-thrust belts (Fig. 1). The western sides of these thrust belts coincide with the Cordillera del Viento (Figs. 1 and 2), where Upper Paleozoic to Mesozoic sedimentary rocks are thrust over Paleogene volcanic rocks (Ramos, 1977; Kozlowski et al., 1996; Zapata et al., 1999). The Cordillera del Viento itself is a narrow range with steep slopes. The 40Ar/39Ar and fission-track ages along with unconformities along the eastern slope indicate that initial uplift of the Cordillera del Viento occurred in the Late Cretaceous (Zapata et al., 1999, 2002; Kay, 2002; Kay et al., this volume, chapter 2). Synorogenic strata are consistent with a series of Eocene and late Miocene reactivations (Ramos, 1998; Zapata et al., 1999; Zapata and Folguera, 2006). To the south of 38ºS, the Main Cordillera was first uplifted in the Eocene, as inferred from fission-track data in the innermost sector of the fold belt (Gräfe et al., 2002). This relief was destroyed by orogenic collapse in the late Oligocene (Ramos and Folguera, 2006; Zapata and Folguera, 2006). Subsequently, this sector was uplifted in the late Neogene, as the Guañacos fold-and-thrust belt initiated through inversion of a series of intra-arc extensional depocenters (Jordan et al., 2001; Radic et al., 2002). STRATIGRAPHIC FRAMEWORK The stratigraphic sequence in the eastern Neuquén Andes begins with Permian volcaniclastic successions that are overlain by the Jurassic to Lower Cretaceous sedimentary sequences of the Neuquén Basin. Magmatic rocks of mainly Cretaceous age lie to the west (Zollner and Amos, 1973; Niemeyer and Muñoz, 1983; De la Cruz and Suárez, 1997; Suárez and Emparán, 1997). Above them are volcaniclastic sequences with ages from 25 Ma to a few thousand years. They can be put in three main groups (Niemeyer and Muñoz, 1983; Suárez and Emparán, 1995). The first main group consists of volcaniclastic sequences with ages between 25 and 15 Ma (Suárez and Emparán, 1995, Jordan et al., 2001; Radic et al., 2002). These rocks are incor-
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porated into the Cura Mallín Formation, which is interpreted to have accumulated in a series of diachronous depocenters in an extensional regime that occupied the intra-arc region from 33° to 46°S (Dalla Salda and Franzese, 1987; Spalletti and Dalla Salda, 1996; Vergara et al., 1997a, 1997b, 1997c; Godoy et al., 1999). Three distinctive units grouped in the Cura Mallín Formation can be mapped in the study area (Fig. 3). The first consists of Oligocene to Miocene nonmarine sedimentary sequences (Sarris, 1964; Zanettini et al., 1987; Leanza et al., 2002) that occur on the easternmost edge of the basin. The other two units occur some 10 km to the west and are composed of the arcderived sediments that comprise the predominant facies of the western slope of the cordillera between 37° and 37°30′S (Niemeyer and Muñoz, 1983). The older of these two, the base of which is unexposed, contains nonmarine white tuffs. The younger consists of nearly 300 m of andesitic breccias associated with irregularly distributed and small basaltic to andesitic lava flows (Fig. 3). Two other two main volcanic groups are composed of latest Miocene to Holocene rocks. The oldest includes volcanic rocks with ages between 5 and 3.5 Ma, which are assigned to the Cola de Zorro Formation (Vergara and Muñoz, 1982; Suárez and Emparán, 1997). They occur in a narrow belt that is locally superimposed on the Upper Oligocene to Lower Miocene volcanic arc. The younger group includes the late Pliocene to early Pleistocene volcanic rocks that underlie the late Pleistocene to Holocene volcanic rocks of the modern arc to the west (Muñoz and Stern, 1988; Muñoz Bravo et al., 1989; Stern, 1989; Lara et al., 2001; Ramos and Folguera, 2006). The eastern slopes of the Andes between 37° and 37°30′S are characterized by these three volcaniclastic sequences (Cura Mallín, Cola de Zorro, and Los Cardos–Centinela Formations), as well as younger stratocones and monogenetic volcanoes and Quaternary landslide, glacial, and fluvial deposits (Figs. 3 and 4). The nearly horizontal successions of the Pliocene Cola de Zorro Formation andesitic lavas, breccias, and nonmarine alluvial fan deposits, which have thicknesses that reach up to 1200 m in this region, are locally accumulated over a basin-wide angular unconformity (Figs. 4 and 5) (Vergara and Muñoz, 1982; Suárez and Emparán, 1997). Among these sequences are Lower Pliocene volcanic rocks between 37°10′ and 20′S, and 70°50′ and 71°10′W that are arranged radially and dip outward with respect to a nearly 15-km-wide, semicircular depression. The depression is interpreted as a moderately eroded volcanic caldera named the Trohunco caldera (Figs. 2 and 3). To the northeast, Upper Pliocene to Lower Quaternary volcanic rocks occur in the amalgamated Los Cardos–Centinela and Palao stratovolcanoes (Figs. 3 and 5). The basal breccias and lavas of the Los Cardos-–Centinela volcano have K-Ar ages of 3.2– 2.6 Ma (Rovere, 1993, 1998; Rovere et al., 2000). Glaciation has been an important agent in modifying the landscape in the region, particularly west of 71°W, where a series of erosive glacial landforms is preserved (Fig. 5) (González Díaz, 1998). At least one episode of glaciation
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Figure 4. Structural cross sections along the Huaraco (A), Lileo (B), and Guañacos (C) transects in the eastern foothills of the Andes between 37º and 37º20′S showing the main structures described in the text. See locations in Figure 3. Detailed diagrams or photos of local structures are in figures indicated by numbers along the transects.
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Figure 5. Pictures of Pliocene to Quaternary volcanic units and glacial deposits related to the Palao and Los Cardos–Centinela (Pu C) volcanic centers (Rovere, 1993). The white lines show the shape of capping ice sheets and the white area, the extent of alpine glaciations.
reshaped the Los Cardos–Centinela and Palao volcanoes (Figs. 5 and 6). During this time, a glacier filled and breached the summit caldera, eroding the volcanic morphology and leaving only the northern and eastern flanks of Los Cardos–Centinela center intact (Figs. 3 and 5). Alpine glacial forms are preserved on the northeastern flank of the volcano and over the extracaldera products (Figs. 5 and 6). Evidence of the more distal effects of glaciation includes remnants of marginal moraines and outwash deposits north of the Huaraco valley. Scattered
pillow-like lavas, which are thought to have formed under ice, also cover the eastern flank of the volcano (Fig. 5). On the lower slopes, an outwash fan, terraced by later fluvial incision of the Huaraco lower valley, is spatially associated with marginal moraines (Fig. 5). Widespread till-like sequences farther west are avalanche deposits associated with the collapse of the western flank of the Los Cardos–Centinela volcano and the southern flank of the Palao volcano (Figs. 3, 5, and 6) (González Díaz et al., 2005a). More
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Figure 6. Map of avalanche deposits associated with the Palao and Los Cardos–Centinela volcanic centers. Note that the Los Cardos avalanche deposits overlie an older glacial morphology. Inset in B shows the Palao fault zone. The fault in the Guañacos fold-and-thrust belt affects both the Paleogene volcanics and the Los Cardos avalanche deposits.
than 50 landslide deposits with volumes greater than 1 × 106 m3 have been identified and mapped in the retroarc area between 36° and 39°S (González Díaz et al., 2000, 2005a, 2005b; González Díaz, 2003; Hermanns et al., 2003). Based on geomorphological evidence and preliminary (36Cl) ages, most of these landslides are assigned to postglacial ages, and a few are considered to be interglacial.
Fifteen kilometers up the Lileo valley, the upper Palao valley has a glacial morphology that is abruptly interrupted by distal avalanche facies from the Los Cardos–Centinela volcano (Figs. 3 and 6) that climb the western side of the Palao valley. Neither glacial erosional features nor glacial material are recognized over this avalanche, suggesting that it has a postglacial age similar to that of other landslides in the region.
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Miocene to Quaternary deformation of the Guañacos fold-and-thrust belt STRUCTURAL TRANSECTS THROUGH THE GUAÑACOS FOLD-AND-THRUST BELT Figure 2 shows the locations of five transects on the eastern slope of the northern Neuquén Andes, and they are described here in order to constrain the age and development of the Guañacos fold-and-thrust belt structure. From north to south, they run through: (1) the Huaraco valley (37°05′–37°10′S) on the flank of the moderately eroded Upper Pliocene to Quaternary Los Cardos–Centinela stratovolcano (Figs. 3 and 4A); (2) the Lileo valley (37°13′S), where the absence of Neogene sequences allows the deformation of Paleogene successions to be accessed (Figs. 3 and 4B); (3) the Guañacos valley (37°16′S), where structures recognized farther north can be seen to deform a Pliocene caldera (Figs. 3 and 4C); (4) the Reñileuvú valley (37°20′S), where subtle changes in the orogenic front geometry and mechanics of deformation can be seen
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(Fig. 3) (Folguera et al., 2004); and (5) the Ñireco valley (37°25′S), where the eastern or inner sectors of the transect are covered by Neogene volcanic rocks that allow the age of the youngest deformation in the area to be assessed. Traverse 1: Structure along the Huaraco Valley The northernmost transect is in the Huaraco valley (Fig. 4A), where the regional structure can be examined relative to the Los Cardos–Centinela and Palao volcanic sequences. Two areas are considered here (Fig. 3): (1) an eastern one, where basaltic and andesitic flows on the eastern slope of the Los Cardos– Centinela volcano are locally affected by east-vergent folds (Fig. 7A), and (2) a western one, where the sequences below the volcano are affected by young and active deformation. Importantly, lavas of two ages occur in the region. The older are the Upper Pliocene to Lower Pleistocene flows that form the
Figure 7. Photos and sketches showing evidence for syngrowth strata in the eastern section of the Guañacos fold-andthrust belt. (A) Low-amplitude anticline in the lower Huaraco valley folding Upper Pliocene to Quaternary lavas and sediments on the eastern slopes of the Los Cardos–Centinela. (B) Syngrowth strata in the Cola de Zorro Formation in the upper Guañacos valley. (C) Quaternary sediments in the lower Guañacos valley associated with fault labeled Fg6 in Figure 4.
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basal part of the eastern half of the Los Cardos–Centinela volcano (Figs. 3, 5, and 6). K-Ar ages from these flows range from 3.2 to 2.6 Ma (Rovere, 1993). The youngest are lavas flows from scattered small Upper Pleistocene cones. Their features fit with eruption under water (or ice), as do those of other Pleistocene synglacial stratovolcanoes in the region (Dixon et al., 1999; Melnick et al., 2003). Similar cones postdate polygenetic centers at the Copahue volcano, 70 km to the south (Fig. 2) (Melnick and Folguera, 2001; Melnick et al., 2006), and at the Chillán volcano, 30 km to the west (Dixon et al., 1999). Western (Outer) Part of the Huaraco Transect. Twelve low-amplitude east-vergent anticlines and synclines occur within a distance of 5 km on the eastern slope of the Los Cardos–Centinela volcano (Figs. 3, 4, and 7A). These folds affect Pliocene lavas, which have a primary dip of 15° to the east (Fig. 8A). A kink increases the average eastern slope of the volcano from 15° to 25°E at a height of 1750 m. The break in slope is coincident with an east-vergent thrust that is associated
with the anticline that exhumes the core and basement of the Los Cardos–Centinela volcano (Fig. 8B). Inner (Eastern) Part of the Huaraco Transect. An anticline with a half-wavelength of 3–4 km is exposed on the southern margin of the Huaraco valley. Its core is formed by Miocene Cura Mallín volcanic sequences and its flanks by the oldest flows from the (late Pliocene?) Los Cardos–Centinela volcano (Fig. 8). The dip of the frontal and eastern limb is slightly steeper (30°E) than the western limb (25°W). This anticline is associated with a 20°W-dipping low-angle reverse fault that thrusts Cura Mallín volcanic rocks over Pleistocene moraine deposits at the bottom of the Huaraco valley (Fig. 9). The dorsal flank of the anticline is unconformably overlain by preglacial Quaternary flows from stratovolcanoes on the northern rim of Los Cardos–Centinela volcano (Fig. 4). To the west, the Palao fault (Fh9 in Fig. 4A) puts Cura Mallín strata over Los Cardos– Centinela rock avalanche deposits that are exposed for a distance of over 8 km (Fig. 6B).
Figure 8. Photos and sketches with east (E) to the left illustrating: (A) The Huaraco anticline that has an angular unconformity between folded Miocene Cura Mallín strata (MCm) below and the units of the Los Cardos–Centinela volcano (Pl/Q C) above. These relations show that the main phase of folding occurred prior to eruption of the Los Cardos–Centinela flows. The sketch below shows the blind fault that this is inferred to control. (B) The Huaraco anticline as exposed in the Guañacos valley further to the south. (C) Close-up of deformed Pleistocene outwash deposits at the frontal part of the reverse fault in the Guañacos valley (see photo in B).
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Traverse 2: Structure along the Lileo Valley
Figure 9. Photos and sketches of reverse fault associated with the forelimb of the Huaraco anticline. The fault puts Lower Miocene volcanic breccias over Pleistocene till deposits that accumulated on the Los Cardos–Centinela crater.
The westernmost part of this transect includes the eastern flank of Cerro Los Tejos (Fig. 3), which is uplifted by a westdipping reverse fault. The hanging wall and footwall of this thrust are better exposed in the Lileo transect discussed in the following paragraph. Evidence for a young age for this thrust comes from the observation that Cura Mallín breccias override postglacial avalanche deposits from the Los Cardos–Centinela and Palao volcanoes (Fig. 10B). Other evidence comes from the Palao valley (Fig. 3), immediately to the east, where avalanche deposits from the Los Cardos–Centinela center are folded into a syncline and anticline (Figs. 3 and 11). Moreover, an east-facing scarp that affects the youngest soil horizon flanks this structure (Figs. 11B and 11C).
Twelve reverse faults that are numbered from Fl1 in the west to Fl12 in the east on Figure 4 form the eastern slope of the Andes in the Lileo transect at 37°13′S (Fig. 3). As in the Huaraco transect, this transect can be divided into two sectors based on deformation style: (1) a western one, where reverse faults and low-amplitude folds affecting mainly sedimentary facies of the Cura Mallín Formation are concentrated in a narrow band, and (2) an eastern one, from the Cerro Los Tejos to the western foothills of the Los Cardos–Centinela volcano, where the volcaniclastic sequences of the Cura Mallín Formation are imbricated into the structures. Outer (Eastern) Part of the Lileo Transect. The eastern part of the Lileo transect includes thrusts Fl1 to Fl6. The two easternmost thrusts (1 and 2) affect lava flows with K-Ar ages of 1.7 ± 0.2 Ma (Folguera et al., 2004), and the consolidated deposits of the Quaternary piedmont zone (Fig. 12A). The easternmost fault (Fl1) has a poor morphological expression. The Los Chacayes (Fl2) fault to the east is actually composed of two reverse faults separated by 30 m. A vertical displacement of 30 m is indicated by the offset between the contact of the lavas and the upper piedmont Quaternary sediments (Fig. 12A). To the north, in the Lileo valley, the western splay of the Los Chacayes fault overrides the Los Miches avalanche deposits (Fig. 3) (Folguera et al., 2004). Fault Fl3 juxtaposes Miocene sedimentary strata over Pliocene volcanic rocks. Fault Fl4 puts sedimentary strata of the Cura Mallín Formation that dip 60°W over Cura Mallín strata that dip 35°W. Similarly, thrust sheet Fl5 puts Cura Mallín strata that dip at 60°W over similar strata that dip at 35°W. Finally, a tight anticline is overridden by fault Fl6. Inner (Western) Part of the Lileo Transect: Cerro Los Tejos, Palao, and Perquiñane-Antiñir Regions. The inner part of the Lileo transect begins with the Pichilenga fault (Fl7, Fig. 4). This fault puts a broad and symmetric syncline with flanks that dip at 25° over a series of footwall folds. A wedge of the Cola de Zorro Formation unconformably overlying the syncline constrains the youngest motion on the fault to be at least late Miocene. The Pichilenga fault (Fl7) is regionally important because it occurs at the main topographic break in the transect and coincides with a major structural boundary in the Oligocene-Miocene Cura Mallín basin. As a result, faults west of the Pichilenga fault involve sedimentary facies of the Cura Mallín Formation, whereas those to the east involve the volcaniclastic facies. Farther west, fault Fl9 juxtaposes a high-amplitude anticline over an anticline, a syncline, and a broad anticline in the Cura Mallín Formation. These footwall folds are in turn cut by the minor Los Rojos reverse fault (Fl8), which can be seen to affect Quaternary fluvial sediments (Fig. 8C). Fault Fl10 to the west puts a tight anticline in the Cura Mallín Formation over thin postglacial avalanche deposits (not shown on Fig. 4) on the southern flank of the Los Cardos–Centinela volcano.
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Figure 10. Photos and sketches (E is east, W is west) of the Los Tejos fault in the Lileo valley illustrating: (A) younger-over-older sequence produced by thrusting within the Cura Mallín Formation. Thrust is labeled F12 in Figure 4. Note that the upper section of the Cura Mallín Formation (Ucm, dark unit in photo) is only exposed on the hanging wall above the Los Tejos fault. This observation is consistent with the Ucm strata forming in the synrift stage of the Cura Mallín basin. Lcm—lower Cura Mallín Formation. (B) Cura Mallín strata thrust over avalanche deposits associated with the collapse of the Los Cardos–Centinela and Palao volcanoes in the Lileo Valley. Sketch below shows faults in more detail. (C) Southern continuation of the Los Tejos fault (Fg12) in the Guañacos valley.
The age of the avalanche deposits could be less than 30,000 yr. Continuing west, the high-angle reverse Palao fault (Fl11), puts Cura Mallín Formation tuffs over avalanche deposits from the western flank of the Los Cardos–Centinela volcano (Fig. 6B). Faults Fl11 and Fl12 are associated with the asymmetric Cerro Los Tejos hill, which has a steep eastern slope that coincides with an active fault. Although details vary from north to south along strike, the overall west-to-east structure of Cerro Los Tejos consists of two tight synclines and an anticline over-
lying a hanging-wall fault. The fault is broken into two reverse faults (Fl11 and Fl12), which override two gentle synclines that are seen in different places along strike (see Fig. 3). These synclines affect the upper levels of the Cura Mallín Formation. Overall, the Los Tejos fault puts upper Cura Mallín Formation volcanic breccias and andesitic lavas over lower Cura Mallín Formation tuffs and ignimbrites, as is the case in a normal fault (Fig. 10A). A logical interpretation is that the younger Cura Mallín sequence preferentially accumulated near the
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Figure 11. Photos of structures associated with the Los Tejos active deformational front. The youngest deformed sediments are avalanche deposits from the Los Cardos–Centinela and Palao volcanoes.
hanging wall of the fault and that the youngest motion on the Los Tejos fault is related to inversion of an extensional depocenter in the eastern part of the Cura Mallín basin. At the hanging wall of the Los Tejos fault, the Cola de Zorro Formation overlies a middle to late Miocene fold (Fig. 4) in the Cura Mallín Formation on an unconformity that dips 5ºW. Traverse 3: Structure along the Guañacos Valley A series of 12 reverse faults occurs in a distance of some 20 km along traverse 3 in the Guañacos valley near 37°14′S (Fig. 4). There is a difference in elevation of ~1000 m between the exposure level of the first fault in the east and the last fault in the west. A distinctive topographic break separates the valley into eastern and western domains. The eastern domain is some 7 km long. The valley in this sector is in a deeply incised postglacial fluvial channel that modifies a glacial valley. The faults in this sector typically exhibit evidence of Quaternary activity. In the western domain, the valley largely retains a Pleistocene morphology, and the thrusts and folds only locally show evidence of Quaternary tectonic activity. Eastern (Outer) Part of the Guañacos Transect. The clearest evidence for Quaternary activity is seen in faults Fg1 to Fg7 in the eastern domain. Beginning in the east, fault Fg1 affects both flat-lying volcanic rocks of probable late Pliocene age as well as the youngest sediments of the region (Figs. 3 and 4).
These sediments are interpreted as a tongue of synorogenic conglomerate and fluvial deposits that accumulated in a depocenter along the footwall of the fault (Folguera et al., 2004). Additional synorogenic sediments along the hanging wall of fault Fg2 accumulated in a gentle syncline that has flanks that dip 6°E and 12°W. To the west, a third reverse fault (Fg3) puts Cura Mallín Formation over synorogenic deposits that accumulated in a westward-thickening wedge east of fault Fg2. Continuing westward, the sedimentary beds in the hanging wall above fault Fg3 dip 30°W until they are overridden by fault Fg4, which places Cura Mallín strata dipping at nearly 60°W above them. No evidence for Quaternary activity was found along the trace of fault Fg4. Farther west, fault Fg5 puts 28°W Cura Mallín strata over Cura Mallín strata and affects Quaternary alluvium in the channel of the Guañacos valley (Fig. 12C). The Quaternary sediments are tectonically repeated at the lowest point of the valley, where a 25°W-dipping sequence includes a layer of cataclastic material. A narrow tongue of alluvium that thickens westward from the fault trace dips at an angle of 25°W near the fault. To the west, 55°W Cura Mallín strata thrusts 28°W Cura Mallín strata (Fg6) offsetting a Quaternary terrace in the Guañacos River (Fig. 12B) and puts Cura Mallín strata over Cura Mallín strata. Less than 600 m to the west, fault Fg7 puts Cura Mallín volcanic rocks that dip at 28°W over Quaternary sediments. Progressive unconformities at the back of this fault can be related to debris talus sediments (Fig. 7C).
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Figure 12. Photos and sketches illustrating deformational style at the orogenic front (E is east). (A) Los Chacayes fault (FI12 in Fig. 4) in the Lileo valley. (B) Fault Fg6 affecting Quaternary sediments in the Guañacos valley. Man is shown for scale. (C) Fault Fg6 affecting Quaternary sediments in the Guañacos valley. See text for more discussion of faults Fg5 and Fg6.
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Miocene to Quaternary deformation of the Guañacos fold-and-thrust belt Fault Fg8 (Fig. 7B) is the easternmost fault that lacks clear evidence of Quaternary activity. Importantly, this fault does not extend into the horizontal volcanic beds that overlie the deformed Cura Mallín Formation above the fault. On a regional scale, the Cura Mallín Formation above the fault dips 45°W before flattening through two kinks to the west. Further, an unconformity separates two post-Pliocene volcanic units over the back limb of the fault. These two volcanic packages have an apex next to the fault trace and rapidly broaden to the west, as would be expected for synorogenic emplacement of the volcanic units (Fig. 7B). The basal package could be Pliocene in age, based on lateral regional correlations, whereas the upper package could be early Quaternary, based on ages of similar volcanic rocks in the vicinity (1.7 Ma) (Folguera et al., 2004). Western (Inner) Part of the Guañacos Transect. Folded Cura Mallín strata are overlain by Pliocene Cola de Zorro strata that are folded into a very broad gentle syncline, which has a half-wavelength of 6 km and flanks that dip 5°W and 7°E (Fig. 13D). Faults Fg10–12 in the inner domain lack evidence of Quaternary deformational activity. To the west, fault Fg10 puts a syncline in the Cura Mallín Formation over the Cola de Zorro Formation (Figs. 4 and 13C). This thrust is notable for affecting the Cola de Zorro Formation and coinciding with an important topographic break. Fault Fg11 uplifts a tight syncline and anticline. These folds are then overridden in <1 km by fault Fg12, which puts a syncline with an 80°W-dipping frontal limb and a gentle back limb above them. Finally, an anticline with a vertical frontal limb and a gentle back limb occurs above the fault Fg12, which is the first thrust in the inner part of the transect that is unconformably overlain by the Cola de Zorro Formation (Fig. 10C). The Pliocene Cola de Zorro strata above fault Fg11 and Fg12 have primary dips that are considered to be related to the only preserved part of the outer rim of the 15-km-wide Trohunco caldera (Figs. 3 and 13A). The Cola de Zorro rocks affected by the folds and thrusts between faults Fg8 and Fg10 are interpreted as the intercaldera facies (Fig. 13B). Their age is considered to be near 3.6 Ma based on dates from volcanic rocks near the Paso de Pichachén (Fig. 3; Muñoz Bravo et al., 1989). Traverse 4: Structure along the Reñileuvú Valley The structures along the Reñileuvú valley transect (Fig. 3) are very similar to those in the more northern traverses. The importance of this traverse is that some structural relationships are particularly well displayed. In one, a thrust fault that is the southern continuation of thrust Fg10 in the Guañacos valley (Fig. 4) puts Miocene ignimbrites over Lower Pliocene volcanic rocks (Fig. 13B). In a second, two synclines are unconformably superimposed (Fig. 13D). This probably shows that some structures were reactivated episodically, first during the middle to late Miocene and then during the last 4–2 (?) m.y. These relations confirm both Miocene and late Pliocene to Quaternary pulses of deformation along the Guañacos fold-andthrust belt.
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Traverse 5: Structure along the Ñireco Valley The traverse along the upper part of the Ñireco Valley (Fig. 2) is important because this valley was filled with postglacial deltaic sediments related to a lake that was damned by landslide deposits (Folguera et al., 2004; González Díaz et al., 2005b). The result was a relatively high valley floor at a high local base level that was only minimally affected by lateral fluvial erosion during postglacial times. As such, younger stratigraphic units that were mostly removed elsewhere in the Guañacos fold-and-thrust belt are well preserved here and provide a marker for young deformation processes. Importantly, the Cola de Zorro Formation can be seen to be gently folded in western sectors of the basin, and a symmetric high-amplitude anticline with limbs of 5º and a NS-trending axis indicate that these strata were affected by late Pliocene to Quaternary deformation near the drainage divide (Fig. 3). DISCUSSION Localization of Quaternary Deformational Activity The full length of the 18-km-long Reñileuvú and the 22-kmlong Guañacos transects along with the eastern two-thirds of the 17-km-long Lileo transect contain faults that show some evidence of late Pliocene to Quaternary activity. This means that nearly half of the eastern slope of the main Andean Cordillera at these latitudes has been affected by young deformation. These observations require a reevaluation of the history of the deformation and uplift of the high cordillera between 36°30′ and 37°30′S, which since the early 1980s has been considered to have essentially ended in the late Miocene, based on the unconformity between the Miocene and Lower Pliocene beds (Niemeyer and Muñoz, 1983; Suárez and Emparán, 1997; Jordan et al., 2001; Radic et al., 2002; Burns, 2002). The observations presented here from the eastern slope show that at least half of the Andes has been built, or at least reshaped, in the Quaternary. The extent of the affected region could be underestimated because the upper levels of the fold-and-thrust belt in the Guañacos, Reñileuvú, and Ñireco transects suggest that the whole eastern flank of the Neuquén Andes has seen some degree of late Pliocene to Quaternary deformation. As such, although the Neogene contractional history of deformation on the eastern flank of the Andes north of 37º30′S began with the Miocene inversion of the Cura Mallín basin, the eastern slope of the Andes needs to be considered as actively deforming and uplifting. Evolution of the Guañacos Fold-and-Thrust Belt between 37° and 37°30’S The structural style of the Guañacos fold-and-thrust belt is one of broad folds that are generally associated with high-angle reverse faults alternating with minor-wavelength folds related to low-angle thrusts. The high-angle faults systematically
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Figure 13. Photos and sketches of the Trohunco caldera region and deformation associated with the Guañacos fold-and-thrust belt. Locations of photos A to D are shown on map. (N is north and W is west). (A) Northern rim of the Trohunco caldera east of the present volcanic front (Sierra Velluda volcano seen to west). (B) Oligocene to Miocene Cura Mallín tuffs overriding Trohunco intracaldera facies at the southern margin of the Reñileuvú valley. (C) Cura Mallín tuffs overriding Lower Pliocene Cola de Zorro Formation volcanic rocks in the Guañacos valley. (D) Broad syncline in Cola de Zorro Formation volcanic rocks overlying tighter syncline in Cura Mallín Formation strata shows evidence for two superimposed deformational phases in the drainage divide between the Guañacos and Reñileuvú valleys.
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Miocene to Quaternary deformation of the Guañacos fold-and-thrust belt involve variable-thickness volcanic units and show youngerover-older relationships. The low-angle thrusts involve sedimentary nonmarine facies, show reverse offsets, and are parts of low-amplitude imbricated fan sheets. It is not certain if the relationship between deformational style and lithology is temporal or spatial or both, because there no outcrops that clearly show the temporal relation between the sedimentary and volcanic facies. The few existing radiometric ages for the volcanic rocks are slightly older than ages inferred from fossils in the sediments (Sarris, 1964; Jordan et al., 2001; Leanza et al., 2002). As such, a plausible model is that the volcaniclastic
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rocks constitute synrift deposits of the Cura Mallín basin, as has been inferred from seismic-reflection lines (Jordan et al., 2001), and that the sedimentary facies are younger sag deposits. In this case, it is logical that sag deposits are preserved on the orogenic front where deformation is incipient and denudation has been minimal (Figs. 3 and 4). Major deformational events in Miocene and Pliocene to Quaternary times, which are inferred from evidence for deformation before and after the Pliocene Cola de Zorro Formation (Fig. 14), could have taken place in different ways. One model is that thin-skinned stacking linked to basement deformation farther
Figure 14. Cross sections showing contrasting deformational styles through the Neuquén Andes. Regions of active deformation in each section are shown in bold. (A) Late Miocene cross section showing deformation related to inversion of the Cura Mallín basin in the western sector and inversion of Lower Jurassic normal faults in the eastern sector. Note the contrasting mechanics of deformation in the region of the Cura Mallín basin, where thick-skinned inversion of basement faults occurred in the west and thin-skinned deformation of sag facies deposits occurred in the east. (B) Late Pliocene to Quaternary cross section showing region where thrusting is occurring (section is close-up of central region). Note that the Agrio and Chos Malal fold-and-thrust belts are now inactive. Contractional deformation is concentrated in the Guañacos fold-and-thrust belt, where the Tertiary Cura Mallín basin is continuing to be inverted. The La Laja strike-slip system to the west becomes active at this time.
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west occurred during the initial development of the fold-andthrust belt and was later overprinted by basement uplifts (Fig. 14). In this case, sag imbricated sequences would have been cannibalized by inversion of preexisting normal faults. This model fits with the observation that the shortening expected for the lowwavelength structures is higher than that associated with the high-angle reverse faults, as would be expected in nonisochronous deformation. Another possibility is that the whole region evolved as a single deformed wedge in which thin-skinned deformation and tectonic inversion were mechanically coupled during both deformational events. In this case, high-angle reverse faults that intersected detachment levels in the sag sequence would produce local thin-skinned deformation. The shortening discrepancy in the first model could be rationalized by invoking a flux of material into the cores of the low-amplitude anticlines, without substantial amounts of shortening. The late Miocene seems to have been a time of widespread contractional deformation north of 37º30′S, as shown by field evidence from the Guañacos and Agrio–Chos Malal fold-andthrust belts. In contrast, the Agrio and Chos Malal fold-and-thrust belts appear to have been fossilized in the late Pliocene to Quaternary, when contractional deformation continued in the inner sectors of the Guañacos fold-and-thrust belt (Fig. 14). A plausible hypothesis for the latest stages of the formation of the western frontal part of the Guañacos fold-and-thrust belt is that it is mechanically coupled with dextral strike-slip displacements in the La Laja fault system described by Melnick et al. (2003; this volume, chapter 4). In this case, basement blocks stacked by transpression at the arc front would have transferred shortening to the Upper Oligocene detachment in the inner retroarc (Fig. 14), explaining its absence beyond this area in the outer fold-and-thrust belt. The absence of younger-than–late Miocene contractional deformations in the Agrio and Chos Malal foldand-thrust belts could be understood as a consequence of being beyond the area of influence of the strike-slip systems that dominate the arc front at these latitudes. CONCLUSIONS The Neogene Guañacos fold-and-thrust belt on the eastern flank of the Neuquén Andes formed in response to inversion of the pre-existing normal faults of the late Oligocene to early Miocene Cura Mallín basin. This stage of compressional deformation occurred at the same time as deformation in the Agrio and Chos Malal fold-and-thrust belts to the east at this latitude. During the late Pliocene to Quaternary, the Guañacos fold-andthrust belt began acting as an out-of-sequence fan of thrusts west of the older, then fossilized, Agrio and Chos Malal foldand-thrust belts. The youthful deformation of the inner sectors of the Guañacos fold-and-thrust belt at these latitudes explains the low altitudes of this region of the Andes compared to the rest. This last phase of basin closure cannibalized the late Pliocene to early Quaternary volcanic arc, and probably earlier thin-skinned structures that developed in the late Miocene. The
inversion of the older extensional structures could be mechanically linked to the La Laja strike-slip system in the intra-arc in the latest phase of deformation. ACKNOWLEDGMENTS This study was made possible by funding from PICT 06729/99 of Agencia Nacional de Promoción Científica y Tecnológica to V.A. Ramos. Reviews by Carlos Costa (Universidad de San Luis) and Matthew Burns (U.S. Geological Survey) helped to improve and clarify concepts in this work. We especially thank Suzanne Mahlburg Kay for discussion and help in the final structure of the paper. REFERENCES CITED Barrientos, S., and Acevedo, P., 1992, Seismological aspects of the 1988–1989 Lonquimay volcanic eruption (Chile): Journal of Volcanology and Geothermal Research, v. 53, p. 73–87, doi: 10.1016/0377-0273(92)90075-O. 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., Jordan, T.E., Copeland, P., and Kelley, S.A., this volume, The case for extensional tectonics in the Oligocene-Miocene Southern Andes as recorded in the 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, doi: 10.1130/2006.2407(08). Dalla Salda, L., and Franzese, J., 1987, Las megaestructuras del Macizo y la Cordillera Norpatagónica Argentina y la génesis de las cuencas volcanosedimentarias Terciarias: Revista Geológica de Chile, v. 31, p. 3–13. De la Cruz, R., and Suárez, M., 1997, El Jurásico de la cuenca Neuquina en Lonquimay, Chile: Formación Nacientes del Bío Bío (38°–39°S): Revista Geológica de Chile, v. 24, no. 1, p. 3–24. Dixon, H., Murphy, M., Sparks, S., Chávez, R., Naranjo, J., Dinkley, P., Young, S., Gilbert, J., and Pringle, M., 1999, The geology of Nevados de Chillán volcano, Chile: Revista Geológica de Chile, v. 26, no. 2, p. 227–253. Folguera, A., and Ramos, V.A., 2002, Partición de la deformación durante el Neógeno en los Andes Patagónicos Septentrionales (37°–46°S): Revista de la Sociedad Geológica de España, v. 15, no. 1–2, p. 81–93. Folguera, A., Araujo, M., Melnick, D., Hermanns, R., García Morabito, E., and Bohm, M., 2003, Seismicity and variations of the crustal tensional state of the retro-arc in the southern Central Andes during the last 5 Ma (37°30′–39°S), in Proceedings of the 10th Congreso Geológico Chileno, Concepción, (CD-ROM). Folguera, A., Ramos, V.A., Hermanns, R., and Naranjo, J., 2004, Neotectonics in the foothills of the southernmost central Andes (37º–38ºS): Evidence of strike-slip displacement along the Antiñir-Copahue fault zone: Tectonics, v. 23, no. TC5008, doi: 10.1029/2003TC011533. Folguera, A., Zapata, T., and Ramos, V.A., 2006, this volume, Late Cenozoic extension and the evolution of the Neuquén 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, doi: 10.1130/2006.2407(12). Godoy, E., Yañez, G., and Vera, E., 1999, Inversion of an Oligocene volcano tectonic basin and uplifting of its superimposed Miocene magmatic arc in the Chilean Central Andes: First seismic and gravity evidences: Tectonophysics, v. 306, p. 217–236, doi: 10.1016/S0040-1951(99)00046-3. González Díaz, E., 1998, Mapa geomorfológico de la hoja geológica 3772-II “Las Ovejas”, Provincia del Neuquén: Buenos Aires, Servicio Nacional de Geología y Minería Argentino, scale 1:250,000, 1 sheet.
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Miocene to Quaternary deformation of the Guañacos fold-and-thrust belt González Díaz, E., 2003, El englazamiento en la región de Caviahue-Copahue: Su reinterpretación: Revista de la Asociación Geológica Argentina, v. 58, no. 3, p. 356–366. González Díaz, E., Fauqué, L., Giaccardi, A., and Costa, C., 2000, Las Lagunas de Varvar Co Campos y Varvar Co Tapia (N. de Neuquén, Argentina): Su relación con avalanchas de rocas: Revista de la Asociación Geológica Argentina, v. 55, no. 3, p. 147–164. González Díaz, E., Folguera, A., and Hermanns, R., 2005a, La avalancha de rocas del cerro Los Cardos (37º10′S, 70º53′O), en la región norte de la Provincia de Neuquén (Argentina): Revista de la Asociación Geológica Argentina, v. 60, no. 1, p. 207–220. González Díaz, E., Folguera, A., and Hermanns, R., 2005b, Reconocimiento y descripción de avalanchas de rocas prehistóricas en el área Neuquina delimitada por los paralelos 37º15′y 37º30′S y los meridianos 70º55′y 71º05′O: Revista de la Asociación Geológica Argentina, v. 60, no. 3, p. 446–460. Gräfe, K., Glodny, J., Seifert, W., Rosenau, M., and Echtler, H., 2002, Apatite fission track thermochronology of granitoids at the south Chilean active continental margin (37º–42ºS): Implications for denudation, tectonics and mass transfer since the Cretaceous, in Proceedings of the 5th International Symposium on Andean Geodynamics: Toulouse, France, IRD Editions,, p. 275–278. Hermanns, R., Folguera, A., and Mardones, M., 2003, Large massive rock slope failures in the Argentine and Chilean Andes between 36 and 38°S: Reno, Geological Society of America–INQUA Abstracts with Programs, v. 16, p. 67. Jordan, T., Burns, W., 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, no. 3, p. 308–324, doi: 10.1029/1999TC001181. Kay, S.M., 2002, Tertiary to Recent transient shallow subduction zones in the Central and Southern Andes, in Proceedings of the 15th Congreso Geológico Argentino, Calafate, Argentina: Actas, v. 3, p. 282–283. Kay, S.M., Burns, W.M., Copeland, P., and Mancilla, O., 2006, this volume, 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, doi: 10.1130/2006.2407(02). Kendrick, E., Bevis, M., Smalley, R., Jr., Cifuentes, O., and Galban, F., 1999, Current rates of convergence across the Central Andes: Estimates from continuous GPS observations: Geophysical Research Letters, v. 26, no. 5, p. 541–544, doi: 10.1029/1999GL900040. Kozlowski, E., Cruz, C., and Sylwan, C., 1996, Geología estructural de la zona de Chos Malal, Cuenca Neuquina, Argentina, in Proceedings, 13th Congreso Geológico Argentino and Congreso de Exploración de Hidrocarburos, Buenos Aires, Argentina: Actas, v. 1, p. 15–26. Lara, L., Rodríguez, C., Moreno, H., and Pérez de Arce, H., 2001, Geocronología K-Ar y geoquímica del volcanismo Plioceno superior– Pleistoceno de los Andes del sur (39°–42°S): Revista Geológica de Chile, v. 28, no. 1, p. 67–90. Lara, L.E., and Folguera, A., 2006, this volume, The Pliocene to Quaternary narrowing of the Southern Andean volcanic arc between 37º and 41ºS latitude, 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, doi: 10.1130/2006.2407(14). Leanza, H., Volkheimer, W., Hugo, C., Melendi, D., and Rovere, E., 2002, Lutitas negras lacustres cercanas al límite Paleógeno-Neógeno en la región noroccidental de la provincia del Neuquén: Evidencias palinológicas: Revista de la Asociación Geológica Argentina, v. 57, no. 3, p. 280–288. Melnick, D., and Folguera, A., 2001, Geología del complejo volcánico Copahue–Caldera Del Agrio, un sistema transtensional activo desde el Plioceno en la transición de los Andes Patagónicos a los Andes Centrales
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(38°S–71°O), in Proceedings, 9th Congreso Geológico Latinoamericano, Symposium “Evolución Tectónica de los Andes,”: Montevideo, Uruguay, University of Montevideo, p. 6–11. Melnick, D., Folguera, A., Echtler, H., Charlet, F., Büttner, O., and De Batist, M., 2003, The Lago del Laja fault system: Active intra-arc orogenic collapse in the southern Central Andes (37°15′S), in Proceedings, 10th Congreso Geológico Chileno, Concepción, (CD-ROM). Melnick, D., Folguera, A., and Ramos, V.A., 2006, Geology of the Copahue volcano–Agrio caldera complex (37º50′S): Structural control, volcanostratigraphy and regional tectonic implications: Journal of South American Earth Sciences (in press). Melnick, D., Rosenau, M., Folguera, A., and Echtler, H., 2006, this volume, Neogene tectonic evolution of the Neuquén Andes western flank (37–39°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, doi: 10.1130/2006.2407(04). Mosquera, A., and Ramos, V.A., 2006, this volume, Intraplate deformation in the Neuquén Embayment, 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, doi: 10.1130/2006.2407(05). Mpodozis, C., and Ramos, V.A., 1989, 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, Circum-Pacific Council for Energy and Mineral Resources: Earth Sciences Series, v. 11, p. 59–90. Muñoz, J., and Stern, C., 1988, The Quaternary volcanic belt of the southern continental margin of South America: Transverse structural and petrochemical variations across the segment between 38° and 39°S: Journal of South American Earth Sciences, v. 1, no. 2, p. 147–161, doi: 10.1016/0895-9811(88)90032-6. Muñoz Bravo, J., Stern, C., Bermúdez, A., Delpino, D., Dobbs, M.F., and Frey, F.A., 1989, El volcanismo Plio-Cuaternario a través de los 38° y 39°S de los Andes: Revista de la Asociación Geológica Argentina, v. 44, p. 270–286. Niemeyer, H., and Muñoz, J., 1983, Geología de la hoja 57 Laguna de La Laja, Región de Bío Bío: Servicio Nacional de Geología y Minería, Santiago de Chile, scale 1:250,000, 1 sheet. Radic, J., Rojas, L., Carpinelli, A., and Zurita, E., 2002, Evolución tectónica de la cuenca Terciaria de Cura Mallín, región cordillerana ChilenoArgentina (36°30′–39°S), in Proceedings, 15th Congreso Geológico Argentino, El Calafate, (CD-ROM). Ramos, V.A., 1977, Estructura de la Provincia de Neuquén, in Rolleri, E.O., eds., Geología y recursos naturales de la Provincia del Neuquén: Buenos Aires, 7th Congreso Geológico Argentino (Neuquén): Asociación Geológica Argentina, Relatorio de la Provincia de Nequén, p. 9–24. Ramos, V.A., 1998, Estructura del sector occidental de la faja plegada y corrida del Agrio, cuenca Neuquina, Argentina, in Proceedings, 10th Congreso Latinoamericano de Geología, Buenos Aires: Actas, v. 2, p. 105–110. Ramos, V.A., and Folguera, A., 2006, Tectonic evolution of the Andes of Neuquén: Constraints derived from the magmatic arc and foreland deformation, in Veiga, G.D., Spalletti, L., Howell, J.A., and Schwarz, E., eds., The Neuquén Basin: A case study in sequence stratigraphy and basin dynamics: Geological Society of London Special Publication 252, p. 15–35. Rovere, E., 1993, K/Ar ages of magmatic rocks and geochemical variations of volcanics from South Andes (37° to 37°15′S–71°W): Japan Volcanological Society Abstracts with Programs, v. 2, p. 107. Rovere, E., 1998, Volcanismo Jurásico, Paleógeno y Neógeno en el noroeste del Neuquén, Argentina, in Proceedings, 10th Congreso Latinoamericano de Geología, Buenos Aires: Actas, v. 1, p. 144–149. Rovere, E., Leanza, H., Hugo, C., Casselli, A., Tourn, S., and Folguera, A., 2000, Hoja geológica Andacollo, Provincia de Neuquén: Buenos Aires, Servicio Nacional de Geología y Minería Argentino, scale 1: 250,000, 1 sheet.
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Vergara, M., Puga, E., Morata, D., Decar, I., Díaz de Federico, A., and Fonseca, E., 1997c, Mineral chemistry of the Oligocene-Miocene volcanism from Linares to Parral, Andean Precordillera, in Proceedings, 8th Congreso Geológico Chileno, Antofagasta: Universidad de Antofagasta, Actas, v. 2, p. 1579–1583. 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., and Folguera, A., 2006, Tectonic evolution of the Andean fold and thrust belt of the southern Neuquén Basin, Argentina, in Veiga, G.D., Spalletti, L., Howell, J.A., and Schwarz, E., eds., The Neuquén Basin: A case study in sequence stratigraphy and basin dynamics: Geological Society of London Special Publication 252, p. 37–56. Zapata, T., Brissón, I., and Dzelalija, F., 1999, The role of the basement in the Andean fold and thrust belt of the Neuquén Basin, Argentina: Thrust Tectonics ’99 Conference Abstracts with Programs, p. 122–124. Zapata, T., Córsico, S., Dzelajica, F., and Zamora, G., 2002, La faja plegada y corrida del Agrio: Análisis estructural y su relación con los estratos Terciarios de la cuenca Neuquina Argentina, in Proceedings, 5th Congreso de Exploración y Desarrollo de Hidrocarburos, Mar del Plata, (CD-ROM). Zollner, W., and Amos, A., 1973, Descripción geológica de la hoja 32b, Chos Malal: Buenos Aires, Carta Geológico Económica de la República Argentina, Boletín no. 143, 91 p., scale 1:200,000, 1 sheet.
MANUSCRIPT ACCEPTED BY THE SOCIETY 22 DECEMBER 2005
Printed in the USA
Geological Society of America Special Paper 407 2006
Late Cenozoic extension and the evolution of the Neuquén Andes Andrés Folguera* Laboratorio de Tectónica Andina, Universidad de Buenos Aires, Consejo Nacional de Investigaciones Científicas y Técnicas, Buenos Aires, Argentina Tomás Zapata Repsol YPF and Universidad de Buenos Aires, Buenos Aires, Argentina Víctor A. Ramos* Laboratorio de Tectónica Andina, Universidad de Buenos Aires, CONICET, Buenos Aires, Argentina
ABSTRACT The eastern slope of the Andes between 36°S and 39°S shows contrasting behavior north and south of 37.5°S to 38°S. The region is notable for a mixed contractional and extensional tectonic regime in the last 5 m.y. that has led to the formation of four broad extensional depocenters and the Guañacos fold-and-thrust belt. The post–late Miocene tectonic evolution of the eastern slope of the Andes south of 37.5°S has been dominated by the development of the Bío Bío–Aluminé and the Loncopué troughs, which have undergone extensional collapse in association with abundant mafic volcanism. The formation of these troughs postdates the middle to late Miocene contractional deformation that inverted the early Miocene Cura Mallín basin. At the same latitude, the forearc has been uplifted by underthrusting and basal accretion. The distribution of late Pliocene to Quaternary deformation along the forearc, arc, and retroarc in this region is typical of a subduction system governed by a negative roll-back velocity. North of 37.5°S, post–late Miocene extensional collapse has occurred in the Las Loicas trough, which extends southeastward from the frontal arc near 35°S to the Tromen region in the retroarc near 37.5°S, and in the Sierra de Reyes trough in the eastern retroarc. The extensional collapse of these basins occurred in a more restricted region than in the troughs south of 37.5°S. The retroarc extension, which is locally accommodated by transtension, postdates late Miocene deformation in the Chos Malal and Malargüe fold-and-thrust belts. To the west, uplift related to shortening has produced the young Guañacos fold-and-thrust belt in the Main Cordillera, which is antithetic to the Benioff zone. This contractional belt is considered to be out-of-sequence in the sense that deformation occurred to the west of the Cretaceous to Miocene fold-and-thrust belt. The structural contrasts north and south of 37.5°S to 38°S are difficult to explain in the context of the modern tectonic setting in which relative convergence parameters, subducting oceanic slab age, and climate are similar. The differences are better explained as responses to contrasting changes in the late Miocene to Holocene geometry of the Benioff zone north and south 37.5°S to 38°S. Keywords: extension, Neuquén Andes, Andean uplift, subduction, slab shallowing, slab steepening. *E-mails:
[email protected];
[email protected] Folguera, A., Zapata, T., and Ramos, V.A., 2006, Late Cenozoic extension and the evolution of the Neuquén 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. 267–285, doi: 10.1130/2006.2407(12). For permission to copy, contact
[email protected]. ©2006 Geological Society of America. All rights reserved.
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INTRODUCTION The south-central Andes (33°–38°S) (Fig. 1) can be considered a classic example of a mountain system formed by subduction of oceanic lithosphere. In a Cretaceous to Miocene framework, deformation in the region has migrated toward the foreland and has been accompanied by a shift of the volcanic front. In the last 5 m.y., the area has seen “normal” subduction, in the sense that arc volcanic activity has been continuous. Over the same time, changes in the dip of the subducting oceanic slab have been postulated to explain the dynamics of the arc and backarc magmatic system between 35° and 38°30′S (Stern, 1989; Kay 2001, 2002; Kay and Mancilla, 2001; Kay et al., this volume, chapter 2). Late Miocene contractional deformation and the eruption of volcanic rocks with arc signatures up to 500 km east of the modern trench have been explained by Miocene shallowing of the subduction slab. Widespread Pliocene to Quaternary alkaline volcanism has been correlated with steeping of the subducting slab. In this paper, the Pliocene
to Holocene structural characteristics and the generally anomalous development of the Andean arc and western retroarc between 36° and 39°S are discussed and shown to be consistent with post-Miocene steepening of the subduction zone. A particularly distinctive feature of the eastern slope of the Andes between 35°S and 39°S is the widespread extension that has occurred in the last 5 m.y. (e.g., Folguera et al., 2003a; Ramos and Folguera, 2006; Zapata and Folguera, 2006). Here, we discuss the evidence for this extension, point out the differences in its character north and south of 37.5°S latitude, and put the extension in a historical perspective relative to the evolution of the Neuquén Basin. Evidence for Pliocene extension in the arc region comes from the 5.6–3 Ma Cola de Zorro basin, which is exposed between 36° and 39°S (Vergara and Muñoz, 1982). The extensional nature of this basin has only recently been recognized (Folguera et al., 2003a; Ramos and Folguera, 2006; Zapata and Folguera, 2006). Evidence for late Pliocene to Quaternary extension in the arc and western retroarc comes from fault-bounded depositional troughs associated with young
Figure 1. Map of the south-central and northern Patagonian Andes on the left shows the location of the study area (white box) relative to the modern Pampean (Chilean) flat-slab, the region where the Nazca plate is subducting at an angle near 30°, the region where late Miocene shallow subduction has been suggested by Kay (2001, 2002), and the principal geologic provinces between 35°S and 38°S. Map on the right shows the principal extensional depocenters and fold-and-thrust belts in the Main Cordillera and western retroarc between 35 and 38°S.
Late Cenozoic extension and the evolution of the Neuquén Andes volcanism (Fig. 1). Those described south of 37.5°S are the Bío Bío–Aluminé (Muñoz and Stern, 1988; García Morabito et al., 2003) and Loncopué (Ramos, 1977) troughs. Those described north of 37.5°S are the Las Loicas and Sierra de Reyes troughs, the characteristics of which are presented here. These troughs developed contemporaneously with the Guañacos fold-andthrust belt between ~36.5°S and 38°S on the eastern slope of the Andes (Fig. 1; Folguera et al., this volume, chapter 11). Contrasts in the tectonic styles north and south of 37.5°S are shown to have a marked influence on the topography of the region and to be compatible with variable amounts of steepening of the subducting plate. SETTING AND PRE-MIOCENE DEFORMATION OF THE NEUQUÉN ANDES The modern Andes between 36° and 39°S are formed by a series of well-defined morphotectonic units with variable degrees of north-to-south expression. From west to east, they are the Coastal Cordillera, the Central Valley, the Main Andes, the retroarc area, and the platform region (Fig. 1A). This region is currently above a segment of the Nazca plate that is subducting at a normal dip (~30°) beneath the South America plate and well south of the shallowing dipping Nazca plate under the modern Pampean flat-slab segment (Fig. 1A). The main emphasis in this paper is on the Main Andes and western retroarc, which are shown in the box in Figure 1B. A broader overview across the entire region can be found in Ramos et al. (this volume, chapter 1) and Kay et al. (this volume, chapter 2). The development of the modern morphotectonic units in the region reflects a series of Cretaceous to Holocene events. The late Miocene to Holocene evolution of the Neuquén Andes has been strongly influenced by the development of the Mesozoic Neuquén Basin, the history of which is summarized in Figure 2 and discussed in more detail in Mosquera and Ramos (this volume, chapter 5). In its early stages, the Neuquén Basin was a Late Triassic to Middle Cretaceous embayment controlled by a series of rifting and sag stages associated with thermal decay (see Vergani et al., 1995). In the late Cretaceous, the Neuquén Basin evolved into a foreland basin east of the Andes (see Barrio, 1990; Tunik, 2001). In the Oligocene to early Miocene, the Neuquén Basin experienced a renewed period of extensional faulting that controlled a series of volcaniclastic depocenters (Fig. 2). The oldest sequences are in the Upper Oligocene to Lower Miocene Cura Mallín Formation. They accumulated in two principal depo-
Figure 2. Stratigraphic chart showing a summary of the major compressional and extensional events in the Andes between 36°S and 39°S. Ages of events are from Ramos (1977, 1998), Vergara and Muñoz (1982), Suárez and Emparán (1995), Jordan et al., (2001), Burns (2002), Gräfe et al. (2002), Kay (2001, 2002), and Kay et al. (this volume, chapter 2).
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centers in a period of no more than 3 m.y. (Jordan et al., 2001; Burns, 2002; Radic et al., 2002; Burns et al., this volume, chapter 8). Seismic-reflection lines show that the thickest sequences are related to synsedimentary extensional structures (Jordan et al., 2001; Burns, 2002). Such extension was a relatively widespread phenomenon in the Southern and Central Andes between 25 and 18 Ma (e.g., Hervé et al., 1993; Suárez and Emparán, 1995; McDonough et al., 1997; Godoy et al., 1999; Muñoz and Araneda, 2000; Rodríguez, 2000; Charrier et al., 2002; Kay and Mpodozis, 2002; Kay and Copeland, this volume, chapter 9). Basin formation continued into the middle to early late Miocene, but none of these basins are clearly tied to extension. One of these basins formed from 15 to 12 Ma on the western side of the Andes south of 38°S (Suárez and Emparán, 1995, 1997). Another is associated with basal volcanic sequences that are respectively dated near 8 Ma in the southern Loncopué trough near Zapala at 39°S (Leanza et al., 2001) and near 9 Ma in the northern Loncopué trough near Andacollo and Las Ovejas at 37°30′S (Pesce, 1981; Burns, 2002). All of these Cretaceous to middle Miocene sequences were affected by contractional deformation some time after 12 Ma and before 4 Ma (Fig. 2). Constraints on the age of this deformation in the Main Cordillera and western slope of the Andes come from Cola de Zorro Formation and other Pliocene volcanic sequences that horizontally overlie deformed sequences (Fig. 2) (Niemeyer and Muñoz, 1983; Suárez and Emparán 1995, 1997; Melnick et al., this volume, chapter 4). Constraints on the age of deformation in the retroarc come from fissiontrack ages in the Cordillera del Viento (Burns, 2002), mammal fossils in synorogenic deposits in the inner Agrio fold-andthrust belt (Zapata et al., 1999, 2002), stratigraphic relations between deformed and undeformed sequences in the western retroarc (Miranda, 1996), and compositional, mineralogical, and geochemical differences in Neogene retroarc magmatic rocks (Kay et al., this volume, chapter 2). Mesozoic to Miocene deformation in the Neuquén Andes occurred in an overall in-sequence manner with respect to compressional inversion of older intra-arc rift faults (Fig. 3). Late Cretaceous to late Miocene deformation was concentrated in the eastern retroarc in the Agrio and Chos Malal fold-and-thrust belts south of 36°S and in the Malargüe fold-and-thrust belt north of 36°S (Fig. 1A; Ramos, 1977; Manceda and Figueroa, 1995; Kozlowski et al., 1996; Zapata et al., 1999). Evidence for a major period of Late Cretaceous exhumation of the Cordillera del Viento region in the western Chos Malal fold-and-thrust belt comes from a 40Ar/39Ar biotite age (Kay, 2002) and fissiontrack data (Burns, 2002). Evidence for Upper Miocene deformation in the Agrio fold-and-thrust belt comes from synorogenic sequences and post-tectonic intrusives (Fig. 3; Ramos and Barbieri, 1989; Zapata et al., 1999, 2002). Miocene deformation was largely accommodated by inversion of Early Jurassic extensional structures and to a lesser extent by thin-skinned deformation (Ramos, 1998).
POST-MIOCENE EVOLUTION OF THE ANDES OF NEUQUÉN AND SOUTHERN MENDOZA PROVINCES An important aspect of the post-Miocene evolution of the Andes between 35°S and 39°S is the differences in extensional and contractional deformation north and south of ~37°30′S. These differences are described in the following sections. Southern Sector: 37°30’S to 39°S Evidence for post-Miocene extension is widespread in the Andes between 37°30′S and 39°S. Neogene to Quaternary normal faults in Chile are discussed by Melnick et al. (2003b; this volume, chapter 4). These faults cut Upper Miocene contractional structures at the northern end of the Liquiñe-Ofqui fault system. They also cut Upper Pleistocene to Holocene sequences in the La Laja and El Barco depocenters (Fig. 3) that are interpreted as transtensional grabens. Farther east, two episodes are recognized in the extensional collapse of the western parts of the fold-and-thrust belt on the eastern slope of the Andes. The first produced the Pliocene Cola de Zorro basin (Vergara and Muñoz, 1982) and the second created the postMiocene Loncopué and Bío Bío–Aluminé basins (Figs. 1 and 3). Support for dynamic extension on the eastern slope of the Andes comes from large negative gravity anomalies (Folguera et al., 2003b) and seismic evidence for crustal attenuation beneath the Loncopué trough (Kind et al., 2002; Yuan et al., this volume, chapter 3). The oldest evidence for post-Miocene extension between 36° and 39°S comes from subhorizontal 5.6 Ma to 3 Ma Cola de Zorro Formation volcanic sequences that overlie folded Upper Miocene rocks on a regional unconformity (Fig. 4) seen on both sides of the Andes (Pesce, 1981; Niemeyer and Muñoz, 1983; Suárez and Emparán, 1997). Evidence for an association with extension comes from wide thickness variations and abrupt thickness changes along faults. Overall, thicknesses are seen to vary from over 1900 m near the Antuco volcano (Vergara and Muñoz, 1982) to a few hundred meters between the Picunleo and Las Damas valley (Figs. 3–5). Progressive unconformities in the Cola de Zorro Formation and younger sequences show that extension was in progress by the Lower Pliocene and continued into the Quaternary (Folguera et al., 2003a). A lack of post-Miocene contractional deformation in the Argentine Andes south of 37°30′S makes extensional control on the distribution of the Lower Pliocene Cola de Zorro volcanic sequences in this area particularly evident. These volcanic rocks can be seen to be distributed along linear features that define quadrangular depocenters covering tens of kilometers. The two major fault systems in the region, the Las Damas and Picunleo, bound extensional depocenters that accumulated sequences with thicknesses reaching more than 1200 m (Fig. 4A). The Las Damas extensional system, which runs through the Las Damas valley, defines a half-graben that extends for more than 20 km along strike (Figs. 3, 4A, and 5).
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Figure 3. Map of the Andean Main Cordillera and adjacent retroarc south of 36°30′S. The map shows the Bio Bio–Aluminé, Loncopué, and Sierra de Reyes troughs, which were the primary regions of extension during the Pliocene to Quaternary. Also shown are the principal Quaternary volcanoes, the Cretaceous to Miocene Chos Malal and Agrio fold-and-thrust belts, the post-Miocene Guañacos fold-and-thrust belt, the Liquiñe-Ofqui fault zone, and the platform area. Radiometric ages of selected volcanic rocks are from Ramos (1981), Muñoz and Stern (1985, 1988), Linares and González (1987), Muñoz Bravo et al., (1989), Rovere (1993, 1998), Vattuone and Latorre (1998), Linares et al. (1999), Kay (2001), Leanza et al. (2001), Burns (2002), and Kay et al. (this volume, chapter 2).
The Picunleo system is related to the Picunleo depocenter (Fig. 3). The 10-km-long Picunleo fault, the northern hanging wall of which collapsed in the early Pliocene, runs through the Río Picunleo valley (Figs. 4A and 5). Evidence for Quaternary extension comes from the Loncopué and Bío Bío–Aluminé troughs (Fig. 3). As recognized by
Burckhardt (1900), these troughs form the two main negative topographic anomalies in this region of the Andes. Both are bordered by N to NNW scarps and linear features that bound central depressions containing monogenetic and polygenetic volcanoes. The Loncopué trough runs along the inner retroarc area between 36°30′and 39°30′S, whereas the Bío Bío–Aluminé
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Figure 4. Maps and photos of the Pliocene to Quaternary transtensional basins and volcanic centers in the inner retroarc as discussed in the text. Locations of maps are shown in a larger context in the Shuttle Radar Topography Mission (SRTM) image in Figure 5. (A) Geological map showing major faults and volcanic centers of the eastern flank of the Andean cordillera. Map (B) and TM satellite image(C) with topographic contours across the Loncopué trough. (D) Cross section across the Agrio caldera and Loncopué trough in the region shown in the map in (C) and the TM image in (D). (E) Photograph showing the horizontal strata in the Cola de Zola Formation and the position of the K-Ar ages mentioned in the text.
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Figure 5. Shuttle Radar Topography Mission (SRTM, 2003) images showing evidence for post-Miocene extension on the eastern flank of the Andean cordillera between 37°S and 38°S. (A) Geological map over shaded digital elevation model (DEM). Ages of the volcanic rocks are from Pesce (1989), Muñoz and Stern (1988), and Linares et al. (1999). Topographic profile shows the relative roughness of the topography across the region. (B–C) Close-ups of regions shown in white boxes in A. Arrows indicate traces of main extensional structures. Image in B shows traces of extensional faults in the Río Picunleo and Las Damas valleys that are discussed in text. Image in C shows a portion of the Loncopué trough.
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trough runs along the arc front between 38°15′ and 39°15′S (Ramos, 1977; García Morabito et al., 2003). The geologic map, topographic images, and cross section in Figures 4B–D, 5A, and 5C show the geologic context, bounding normal faults, and stratigraphic units of the Loncopué trough. The volcanic deposits at the base of the Loncopué trough are dated at 9 Ma at the northern extreme of the trough (Burns, 2002) and at 8.6 Ma (Linares and González, 1987; Leanza et al., 2001) near the southern extreme. Other radiometric ages for Loncopué trough volcanic rocks plotted in Figure 3 are 5 Ma (Linares et al., 1999), 2.3 Ma (Leanza et al., 2001), 0.47 Ma (Linares and González, 1987), and 1.6–0.6 Ma (Linares et al., 1999). The age of the Bío Bío–Aluminé basin is less constrained. An age of 3.4 Ma from Vattuone and Latorre (1998) is shown on Figure 3 for the volcanic rocks on the western side of the trough. The volcanic rocks associated with both troughs have mafic alkaline compositions and low 87Sr/86Sr isotopic ratios (Muñoz Bravo et al., 1989; Vattuone and Latorre, 1998). Northern Sector: 35°S to 37°30’S In contrast to the southern sector, the arc and western retroarc regions north of 38°S are characterized by high-angle late Miocene to Quaternary east-vergent reverse faults, which cut sequences that were previously uplifted on Cretaceous to Miocene thrusts. This fan of reverse faults is called the Guañacos fold-and-thrust belt and is described in detail (Figs. 1, 3, 6A, and 6B) by Folguera et al. (this volume, chapter 11). The western (inner) part of the Guañacos belt is linked with tectonic inversion of the early Miocene Cura Mallín basin (Burns, 2002). The faults in Chile and near the drainage divide were active in the late Miocene (Niemeyer and Muñoz, 1983; Melnick et al., this volume, chapter 4). Those in the eastern (outer) part of the belt affect Pliocene and Quaternary rocks in the northernmost part of the Loncopué trough (Figs. 3, 6A, and 6B) (Folguera et al., 2004; Folguera et al., this volume, chapter 11). Farther east, the Chos Malal and northern Agrio fold-andthrust belts have a complex history (Figs. 3 and 6). Episodes of Late Cretaceous and Miocene contraction advanced through the area inverting older extensional structures in an in-sequence progression. This evolution was interrupted by a short period of quiescence, which was followed by local extensional collapse of the late Miocene uplifts. The Sierra de Reyes trough in the eastern part of the Chos Malal belt formed in this period (Figs. 3 and 6A). The Sierra de Reyes trough evolved synchronously with the eruption of the youngest volcanic rocks seen in the Chos Malal and northern Agrio belts (Figs. 6A, 6C, and 7). These volcanic rocks consist of isolated basaltic domes and lava flows. Their ages range from 3.2 to 1.5 Ma (Linares and Gonzalez, 1987), and their eruptive centers are considered to be related to extensional faults (Bermúdez and Delpino, 1989). One of these centers is the Cochiquito volcano (Bermúdez, 1985) in the northern
part of the Sierra de Reyes trough (Fig. 7B). New seismic data (Fig. 7C) and borehole information (Fig. 7D) show that the Cochiquito flows erupted along a normally reactivated reverse fault in the Chos Malal belt (Figs. 7B and 7E). As seen in Figures 7C and 7D, the structure under the Cochiquito volcano is largely controlled by a west-dipping basement-cored reverse fault called the Sierra de Reyes fault. Reverse motion on this fault is interpreted to be associated with the uplift of the Mesozoic basement block that forms the Sierra de Reyes (Figs. 6 and 7). The seismic line in Figure 7C shows evidence for normal motion on this fault, which is consistent with a relation between extension and the eruption of the Cochiquito volcano. Other evidence for normal reactivation of the Sierra de Reyes thrust could come from the linear northsouth course of the Río Grande River in this region (Figs. 1 and 7E). The eruption of young basaltic cones (Fig. 7E) in the core and on the flanks of the Sierra de Reyes, basaltic cones on the Chihuidos high (Fig. 3) to the south (Ramos, 1977, 1981), and extensional scarps on the eastern side of the Agrio fold-andthrust belt are all consistent with Pliocene extension. The Loicas trough to the east and north is a NNW-trending depression associated with abundant mafic and silicic volcanic centers (Figs. 1 and 8). The physical expression of this trough runs for more than 250 km. The northern end of the trough near 35°S intersects the volcanic arc at the Planchón-Azufre caldera. The southern part of the trough near 37°30′S includes the Tromen volcano, which is more than 100 km into the retroarc. The Las Loicas trough can be divided into a northern and a southern part based on structural features and volcanic rock types. The northern part of the trough coincides with the Río Grande valley, which is interpreted as a half-graben, where its eastern side is bounded by a west-dipping normal fault (Fig. 8). The western margin of the trough in this sector runs close to the Southern volcanic zone arc front. The trough itself is partially filled with andesitic lava flows and ignimbritic deposits from volcanic centers in the western half of the trough. The three major volcanic sources in the north are the Planchón-Azufre caldera, the Calabozos caldera, and the Puelche volcanic field. Those farther south are the Bobadilla, Mary, and Varvarcó calderas. Petrological and geochemical studies of the northern centers by Hildreth et al. (1984, 1991, 1999), González Ferrán (1995), and others show that the silicic volcanic rocks have crustal affinities. Hildreth et al. (1999) argued that the silicic volcanism was genetically related with a change from contractional to extensional conditions in the retroarc. Contractional deformation in this region is well documented until the late Miocene (Manceda and Figueroa, 1993, 1995; Ramos et al., 1996), when it was replaced by an extensional or neutral regime that would favor underplating of mafic magmas and crustal melting. Extensional faults affecting late Pliocene to Quaternary volcanic rocks (see Ramos et al., this volume, chapter 1) support the proposal of Hildreth et al. (1999).
Figure 6. (A) Regional cross section from the arc into the retroarc near 37°S. The western part of the section shows the out-of-sequence Guañacos fold-andthrust belt that marks the Pliocene-Quaternary orogenic front, and the eastern part shows the inactive Cretaceous to Miocene Chos Malal fold-and-thrust belt. The cross section of the Chos Malal belt is from Zapata et al. (2002). (B) Shuttle Radar Topography Mission (SRTM) image across the region of the cross section in A. Radiometric ages for the volcanic rocks in the region are from Rovere (1993, 1998). (C) Close-up shows the region of the Centinela volcanic center (white box in B) at the northern end of the Loncopué trough. Ages are from Linares and Gonzalez (1987).
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Figure 7. Images of the Sierra de Reyes trough region showing evidence for Pliocene extension in the retroarc. (A) Shuttle Radar Topography Mission (SRTM) image shows Sierra de Reyes relative to the Cordillera del Viento and location of cross sections, maps, and images in parts B to E. (B) SRTM image shows location of extensional fault (indicated by arrows) and location of Cochiquito volcano. (C) Cross section based on seismic image and borehole in D. (D) The section shows a basement-cored fault that could be linked to the Miocene uplift of the Sierra de Reyes. This same structure is interpreted to have been reactivated after 3 Ma by the extensional faulting that controlled the emplacement of the Cochiquito volcano. (E) Geologic map of area in box in B showing regional distribution of faults and mafic volcanic rocks.
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Figure 8. Map shows the Las Loicas trough running southeastward from the arc in central Mendoza to the retroarc in the northern part of Neuquén Province. A comparison with Figure 1 shows that the trough crosses the western Chos Malal and Malargüe fold-and-thrust belts. Important features to note are the abundant volcanic rocks in the trough, the correspondence of the inflection of the arc front near 35°S with the northern part of the trough, and the coincidence of the trough with the region of Miocene shallow subduction proposed by Kay (2001) (see Fig. 1). The distribution of volcanic centers and their ages are based on information in Hildreth et al. (1984, 1991, 1999), González Ferrán (1995), Miranda (1996), and Kay et al. (this volume, chapter 2).
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The southern part of the Las Loicas trough is characterized by a series of normal faults that affect late Miocene to Quaternary volcanic rocks in the retroarc. These normal faults, which show polarity jumps, are associated with discrete transfer faults and broad folds that define slip-accommodation zones (Fig. 8). The principal extensional faults are located in the regions of Pliocene silicic calderas to the west (Bobadilla, Varvarcó, Domuyo, de la Cruzada), and the Waile and Negro de Tromen volcanoes and surrounding centers to the east (Fig. 8). The evolution of the Pliocene to Holocene volcanic rocks from the Waile and Tromen and nearby mafic centers and the silicic rocks from the Cerros Tilhué and Bayo centers are discussed by Llambías et al. (1982) and Kay et al. (this volume, chapter 2). TOPOGRAPHIC DIFFERENCES NORTH AND SOUTH OF 37.5°S The contrasting structural styles in the Neuquén and southern Mendoza Andes in the regions north and south of 37.5°S are expressed in topographic images and profiles in Figures 9A to 9C and in the three dimensional perspective diagram in Figure 10. The main differences can be seen in transects north and south of 37°30′S. Transect A at 37°15′S crosses the Andes at the latitude where contractional structures in the Guañacos belt deform Pliocene to Quaternary sequences. Transect B at 37°45′S crosses the Andes where extension has played a major role in defining the morphology of the arc and western retroarc. The main west-to-east features in these transects are discussed in the following paragraphs. Starting in the west, the Coastal Cordillera in Chile (Fig. 9B) is formed by Upper Paleozoic rocks (Hervé et al., 1988). North of 37°30′S, the Coastal Cordillera is a relatively low area with little relief that produces a small effect on the western end of profile A in Figure 9C. In contrast, south of 37°30′S, where Paleozoic units in the Coastal Cordillera are being exhumed by the stacking of oceanic slivers and accretionary sediments on high-angle reverse faults (Melnick et al., 2003a), a relatively high and moderately rough topography is seen on the western end of profile B (Fig. 9C). Farther east, the topography of the central depression, named the Central Valley of Chile (Fig. 9), shows little expression on either transect. The Central Valley is thought to have been a locus of extension and subsidence during the late Oligocene to early Miocene (McDonough et al., 1997; Vergara et al., 1997). Continuing east, the elevation of the Andean Main Cordillera between 36°S and 39°S is generally lower then in the cordillera to the north and generally similar to the cordillera to the south (Fig. 1). The contrast to the north reflects differing times, amounts, and mechanisms of uplift because the Andes between 33°S and 35°S was the locus of active deformation and denudation related to late Miocene to Pliocene contraction and uplift (e.g., Ramos et al., 1996; Kay et al., 2005). Evidence for uplift comes from 40Ar/39Ar ages of plutons (Kurtz et al., 1997) and fission-track analyses in the western flank of the chain
(Maksaev et al., 2003). In contrast, fission-track analyses in the northern Patagonian Andes indicate that the axial part of the Andes south of 40°S was denuded in the Pliocene to Quaternary time (Gräfe et al., 2002), in accord with the structural and neotectonic study of Lavenu and Cembrano (1999). Between 36° and 39°S, the axial part of the Andes north of 37°30′S is higher than to the south (Figs. 1 and 9). Rocks exposed north of 37°30′S are largely Paleogene in age (Fig. 5; Burns, 2002; Radic et al., 2002; Folguera et al., 2004). These exposures occur on the eastern side of the Andes along the Copahue-Antiñir fault zone (Figs. 3 and 10), which bounds the eastern front of the late Miocene to Quaternary Guañacos foldand-thrust belt (Folguera et al., 2004). The Copahue-Antiñir fault is considered to be the northern continuation of the Liquiñe-Ofqui fault zone (Fig. 1 and 3). In the topographic profile in transect A at 37°15′S (Fig. 9C), the western part of the main Andes has a rough topography where intra-arc rifting is occurring in the La Laja lake region. In contrast, a smoother profile is seen to the east in the Guañacos fold-and-thrust belt, where the maximum orogenic relief occurs. To the south, the rocks exposed along the axial parts of the Andes between 37°30′S and 39°S are principally Pliocene to Quaternary in age (Vergara and Muñoz, 1982; Suárez and Emparán, 1997; Melnick et al., 2002). The basement has relatively older denudation ages (Gräfe et al., 2002), and evidence for substantial Pliocene-Quaternary uplift is lacking (Folguera et al., 2004). The region from 37.5°S to 38°S coincides with the area where the Liquiñe-Ofqui fault zone runs southeast through the Upper Pliocene to Lower Quaternary arc volcanoes east of the drainage divide (Miranda et al., this volume, chapter 13). South of 38°S, the Liquiñe-Ofqui fault coincides with the Upper Pleistocene to Holocene arc front (see Fig. 3; Lavenu and Cembrano, 1999; Cembrano et al., 2000). Fission-track analyses between 38°S and 40°S show that the uplift of the axial part of the intra-arc zone is as old as Eocene (Gräfe et al., 2002). Uplift in the retroarc to the east is considered to be late Miocene in age (Zapata et al., 1999, 2002; Burns, 2002). The topographic profile for transect B near 37°45′S (Fig. 9C) shows a rough topography across the whole main Andes where neotectonics is occurring, and a lower-wavelength topography with peaks related to volcanic centers in the retroarc to the east. In summary in the region between 36° and 39°S, uplifted areas with broad and smooth topographic features closely correlate with regions of contractional deformation, whereas high areas with irregular topography correspond to regions of recent or active extension. Regions of contractional deformation include: (1) the Guañacos fold-and-thrust belt in profile A where peneplains have recently or are actively uplifting in association with contractional deformation (Folguera et al., 2004), and (2) the forearc in profile B in the region of active basal accretion and uplift. The regions of extensional deformation coincide with the intra-arc in profile A and the intra-arc and western retroarc areas in profile B (Folguera and Ramos, 2000; Melnick et al., 2002; Ramos and Folguera, 2006; Zapata and Folguera, 2006).
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Figure 9. Topographic image, map, and topographic profiles from the Coastal Ranges in the forearc across the Andean Main Cordillera into the western retroarc between 37°S and 38°S. (A) Digital elevation model (DEM) based on Shuttle Radar Topography Mission (SRTM, 2003) data. (B) Map showing regions discussed in text where contractional and extensional tectonics have occurred. (C) Topographic profiles showing relief patterns along transects A and B.
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Figure 10. Perspective three-dimensional block diagram showing distribution of regions where extension has occurred (horizontal lined patterns), the sense of slip, and regions where contraction has occurred in the last 15 m.y. (in light gray). Upper left inset shows location of block diagram relative to the Neuquén Embayment.
Late Cenozoic extension and the evolution of the Neuquén Andes DISCUSSION: ORIGIN OF TECTONIC DIFFERENCES The structural and topographic transects between 37° and 38°S show that recent tectonic activity across this area is characterized by extensional depocenters and narrow zones with contractional structures. The topography reflects the stress regimes that have been active in the last 5 m.y. Where extensional structures cut previous contractional structures, the topographic pattern shows positive areas with a “basin and range” pattern (Fig. 9). These “basin and range” patterns occur at the highest parts of the cordillera and are restricted to a 30-km-wide strip near 37°S along the Quaternary volcanic front, and a zone from the volcanic front to an inner retroarc zone trough near 38°S, some 90 km to the south. Where recent contractional structures dominate, broad positive and smooth features with wavelengths that vary between 60 and 50 km characterize the relief. Two areas show this pattern, a northern area from the volcanic arc front to the inner retroarc area along transect A north of 37°30′S, and a southern area in the forearc region along transect B south of 37°30′S (Fig. 9). Although these areas in the Guañacos fold-and-thrust belt and the Coastal Cordillera exhibit similar topographic patterns, they have been formed, respectively, by retroarc-antithetic or forearc synthetic development in relation to the Benioff zone fan of thrusts (Figs. 9 and 11). The topographic and structural differences seen in the transects between 37°S and 38°S are not clearly related to contrasting convergence vectors between the South American and Nazca plates (Pardo Casas and Molnar, 1987; Somoza, 1998), changes in the present Wadati-Benioff zone geometries along the continental margin, or discontinuities in the border geometry. In contrast, over the last 10 m.y., different paleo-Benioff configurations in these regions have been suggested based on analyses of arc and retroarc dynamics (Kay, 2001, 2002; Kay et al., this volume, chapter 2; Ramos and Folguera, 2006). South of 38°S, the volcanic arc progressed eastward from 8 Ma until an important westward retreat occurred around 6–5 Ma. This retreat coincides with the inception of the extensional regime that is evidenced by the abrupt thickness changes in the Pliocene sequences in the Main Andes and inner retroarc (Folguera et al., 2003a; Zapata and Folguera, 2006; Ramos and Folguera, 2006). North of 37°30′S, a westward narrowing of subduction-influenced volcanism in the Pliocene (Kay, 2001, 2002; Kay et al., this volume, chapter 2) correlates with the retroarc extension that began in the Pliocene and the formation of the antithetic Guañacos fold-and-thrust belt on the eastern flank of the Main Cordillera. A model calling for differing degrees of negative slab rollback associated with contrasting amounts of steepening of the subduction zone north and south of 38°S in the last 5 m.y. can explain many of these observations. The model is shown in Figure 11 for profiles near 37°–38°S and 38°–39°S. In the northern profile (Fig. 11A), steepening is shown as being more pronounced because the Miocene slab dip has been argued to have
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been the shallowest in this region. The dip change under the retroarc would have been more than 10° and perhaps as much as 30°. In the southern profile (Fig. 11B), the steepening of the slab is shown as being associated with a change in slab dip from ~30° to 40° or about a 10° change under the retroarc. These changes in slab dip along with a retreating plate boundary can be reconciled with: (1) an asthenospheric influx into a thickening mantle wedge, (2) widespread extension in the arc and retroarc contemporaneous with subduction, (3) weakening of the lower continental lithosphere leading to a relatively low topographic relief, preservation of young stratigraphic sequences, and accumulation of only moderate amounts of synorogenic deposits in the inner retroarc, (4) lateral horizontal collapse beneath the arc platform, (5) creation of a fan of out-ofsequence thrusts as the retroarc deformed antithetically to the Benioff zone, and (6) development of a synthetic thrust in the forearc, in relation to the Wadati-Benioff zone. Another question is why Pliocene-Quaternary volcanic rocks in the Las Loicas trough are dominated by silicic volcanic rocks, whereas those in the Bío Bío–Aluminé and the Loncopué troughs are largely mafic volcanic rocks. An explanation could be related to larger amounts of late Miocene crustal thickening in the Malargüe fold-and-thrust belt in the north than in the Agrio and Chos Malal fold-and-thrust belts in the south (Fig. 1). More importantly, mantle upwelling related to a greater steepening of the oceanic slab could enhance crustal melting in a thicker crust. CONCLUSIONS Extension has differentially affected the arc and western retroarc area of the Neuquén Andes since the late Miocene. South of 38°S, the whole arc and western retroarc has been attenuated, first between 5 and 3 Ma, and then during the last 2 m.y. The forearc has been a locus of active uplift in response to crustal thickening. North of 38°S, extension seems limited to the period between 5 and 3 Ma, and the antithetic Guañacos fold-and-thrust belt formed on the eastern margin of the Andes. The differential behavior north and south of 37°30′S to 38°S can be correlated with the changes in Benioff-zone geometry at 7 and 5 Ma along the Andean margin at these latitudes (Kay, 2001, 2002; Kay et al., this volume, chapter 2). Extension can be correlated with steepening of the slab during this period, where the whole region was affected to varying degrees. In contrast, retroarc contraction is present only in the north. We envision that asthenospheric influx caused by broadening of the asthenospheric wedge during steepening would have interacted with a hydrated lower lithosphere and created a weak crust in that region (Fig. 11). Upon a return to a normal subduction regime, this crust would have yielded and produced the shortening seen in the Guañacos fold-and-thrust belt. The present segmentation of the Andes between 36° and 39°S can thus be attributed to postMiocene changes in the subduction-zone geometry and the extensional regime documented through the region.
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Figure 11. Lithospheric-scale cartoon cross sections approximately along transects A and B in Figures 9 and 10 showing contrasting geology and the effects of negative trench roll-back. The gray lines in the asthenosphere show the inferred shape of the Wadati-Benioff zone at 5 Ma, and the black lines, the inferred shape at present. Solid dots in profile B are earthquake hypocenters from Bohm et al. (2002). Evidence for a retreating plate boundary since the late Miocene comes from: (1) retroarc extension contemporaneous with subduction shown in both profiles, (2) relatively low elevations in both profiles, (3) relatively little erosion as evidenced by exposures of young stratigraphic sequences in both profiles—older sequences occur in profile A where east-vergent Guañacos thrust system occurs and active denudation is taking place in the retroarc, (4) accumulation of a moderate amount of synorogenic deposits in the inner retroarc in profile B, (5) the development synthetic to the Wadati-Benioff zone of a thrust system in the forearc in profile B, and (6) the development antithetic to the Wadati-Benioff zone of a fan of thrusts in profile A (Guañacos fold-and-thrust belt) where the slab-dip change is inferred have been the greatest.
Late Cenozoic extension and the evolution of the Neuquén Andes ACKNOWLEDGMENTS This study was made possible by funding from PICT 06729/99 of Agencia Nacional de Promoción Científica y Tecnológica to V.A. Ramos. We thank Suzanne Mahlburg Kay and Ben Brooks for reviews and suggestions that substantially improved the presentation and clarified the concepts in this paper. REFERENCES CITED Barrio, C., 1990, Late Cretaceous–Early Tertiary sedimentation in a semi-arid foreland basin (Neuquén Basin, western Argentina): Sedimentary Geology, v. 66, no. 3–4, p. 255–275, doi: 10.1016/0037-0738(90)90063-Y. Bermúdez, A., 1985, Los basaltos post-Pliocenos entre los 36° y 37° de latitud sur, Provincia de Mendoza, Argentina, in Proceedings, 6th Congreso Geológico Chileno: Antofagasta, Universidad del Norte, v. 4, p. 52–67. Bermúdez, A., and Delpino, D., 1989, La provincia basáltica Andino cuyana (35°–37° L. S): Revista de la Asociación Geológica Argentina, v. 44, no. 1–4, p. 35–55. Bohm, M., Lüth, S., Echtler, H., Asch, G., Bataille, K., Bruhn, C., Rietbrock, A., and Wigger, P., 2002, The Southern Andes between 36° and 40°S latitude: Seismicity and average seismic velocities: Tectonophysics, v. 356, p. 275–289, doi: 10.1016/S0040-1951(02)00399-2. Burckhardt, C., 1900, Coupé geologiqué de la cordillére entre Las Lajas et Curacautín: Museo de la Plata, Anales, Sección Geológica y Mineralógica, 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., Jordan, T.E., Copeland, P., and Kelley, S.A., this volume, The case for extensional tectonics in the Oligocene-Miocene Southern Andes as recorded in the 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, doi: 10.1130/2006.2407(08). Cembrano, J., Schermer, E., Lavenu, A., and Sanhueza, A., 2000, Contrasting nature of deformation along an intra-arc shear zone, Liquiñe-Ofqui fault zone, southern Chilean Andes: Tectonophysics, v. 319, p. 129–149, doi: 10.1016/S0040-1951(99)00321-2. Charrier, R., Baeza, O., Elgueta, S., Flynn, J., Gans, P., Kay, S., Muñoz, N., Wyss, A., 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°SL): Journal of South American Earth Sciences, v. 15, p. 117–139, doi: 10.1016/S0895-9811(02)00009-3. Folguera, A., and Ramos, V.A., 2000, Control estructural del Volcán Copahue: Implicancias tectónicas para el arco volcánico Cuaternario (36°–39°S): Revista de la Asociación Geológica Argentina, v. 55, p. 229–244. Folguera, A., Ramos, V.A., and Melnick, D., 2003a, Recurrencia en el desarrollo de cuencas de intraarco, Cordillera Neuquina (37°30′): Revista de la Asociación Geológica Argentina, v. 58, no. 1, p. 3–19. Folguera, A., Introcaso, A., and Ramos, V.A., 2003b, Atenuamiento cortical y extensión activos en el arco y retroarco Andino entre los 37°–39°S a partir de estudios gravimétricos y geológicos de superficie, in Proceedings, 10th Congreso Geológico de Chile: Concepción, Electronic files, (CD-ROM). Folguera, A., and Ramos, V.A., Hermanns, R., and Naranjo, J., 2004, Neotectonics in the foothills of the southernmost central Andes (37°–38°S): Evidence of strike-slip displacement along the Antiñir-Copahue fault zone: Tectonics, v. 23, TC5008, doi: 10.1029/2003TC011533. Folguera, A., Ramos, V.A., González Díaz, E.F., and Hermanns, R., 2006, this volume, Miocene to Quaternary deformation of the Guañacos fold-andthrust 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
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Late Cenozoic extension and the evolution of the Neuquén Andes Ramos, V.A., 1998, Estructura del sector occidental de la faja plegada y corrida del Agrio, cuenca Neuquina, Argentina, in Proceedings, 10th Congreso Latinoamericano de Geología: Buenos Aires, Servicio Geológico Argentino, Actas, v. 2, p. 105–110. Ramos, V.A., and Folguera, A., 2006, Tectonic evolution of the Andes of Neuquén: Constraints derived from the magmatic arc and foreland deformation, in Veiga, G.D., Spalletti, L., Howell, J.A., and Schwarz, E., eds., The Neuquén Basin: A case study in sequence stratigraphy and basin dynamics: Geological Society of London Special Publication 252, p. 15–35. Ramos, V.A., Cegarra, M., and Cristallini, E., 1996, Cenozoic tectonics of the high Andes of the west-central Argentina (30°–36°30′S S): Tectonophysics, v. 259, p. 185–200, doi: 10.1016/0040-1951(95)00064-X. Ramos, V.A., and Kay, S.M., 2006, this volume, Overview of the tectonic evolution of the southern Central Andes of Mendoza and Neuquén (35°–39°S latitude), 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 latitude): Geological Society of America Special Paper 407, doi: 10.1130/2006.2407(01). Rodríguez, J., 2000, Reconstrucción de la evolución de la Cuenca del Golfo de San Jorge durante el Terciario, in Proceedings, 2nd Congreso Latinoamericano de Sedimentología: Mar del Plata, Instituto Argentino del Petróleo y del Gas. Rovere, E., 1993, K/Ar ages of magmatic rocks and geochemical variations of volcanics from South Andes (37° to 37°15′S–71°W), in Proceedings, 2nd Japan Volcanological Society: p. 107. Rovere, E., 1998, Volcanismo Jurásico, Paleógeno y Neógeno en el noroeste del Neuquén, Argentina, in Proceedings, 10th Congreso Latinoamericano de Geología: Buenos Aires, Servicio Geológico Argentino, v. 1, p. 144–149. Somoza, R., 1998, Updated Nazca (Farallon)–South American relative motions during the last 40 my: Implications for mountain building in the Andes: Journal of South American Earth Sciences, v. 11, p. 211–215, doi: 10.1016/S0895-9811(98)00012-1. Stern, C., 1989, Pliocene to Present migration of the volcanic front, Andean Southern volcanic front: Revista Geológica de Chile, v. 16, no. 2, p. 145–162. STRM, 2003, Shuttle Radar Topography Mission 90m-dataset. NASA-NIMA: http://srtm.usgs.gov/. Suárez, M., and Emparán, 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.
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Geological Society of America Special Papers Upper Pliocene to Lower Pleistocene volcanic complexes and Upper Neogene deformation in the south-central Andes (36°30 ′−38°S) Fernando Miranda, Andrés Folguera, Pablo R. Leal, José A. Naranjo and Abel Pesce Geological Society of America Special Papers 2006;407;287-298 doi: 10.1130/2006.2407(13)
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Geological Society of America Special Paper 407 2006
Upper Pliocene to Lower Pleistocene volcanic complexes and Upper Neogene deformation in the south-central Andes (36°30 ’–38 °S) Fernando Miranda* Servicio Geológico Minero Argentino, Julio A. Roca 651, Capital Federal, Argentina Andrés Folguera Laboratorio de Tectónica Andina, Universidad de Buenos Aires, Pabellón 2 Ciudad Universitaria, Capital Federal, Argentina Pablo R. Leal Cátedra de Mineralogía, Universidad de Buenos Aires, Pabellón 2 Ciudad Universitaria, Capital Federal, Argentina José A. Naranjo Servicio Nacional de Geología y Minería de Chile (SERNAGEOMIN, Chile), Av. Santa María 0104, Providencia, Santiago, Chile Abel Pesce Servicio Geológico Minero Argentino, Area de Geotermia, Julio A. Roca 651, Capital Federal, Argentina
ABSTRACT The Agrio (37°51’S, 70°26’W), Vilú Mallín (37°28’S, 70°45’W), Trohunco (37°18’S, 71°01’W), Domuyo (36°38’S, 70°26’W), and Los Cardos–Centinela (37°06’S, 70°52’W) volcanic complexes in Argentina are the principal Upper Pliocene to Lower Pleistocene volcanic complexes east of the Andean Main Cordillera and the modern Southern volcanic zone arc front. These complexes are part of the Upper Pliocene to Lower Pleistocene volcanic arc that was on the eastern flank of the Andes at that time. The volcanic rocks provide constraints on the age and style of Neogene deformation in the modern backarc between 36°30’ and 38°S. New and published K-Ar ages along with stratigraphic and structural relations show that the region was affected by a late Miocene compressional deformation between 9 and 6.8 Ma. A more heterogeneous picture emerges for younger deformation in the region. The most important structures include a N-NW–trending contractional fault system that connects the Trohunco and Los Cardos–Centinela complex, and a NE-trending extensional fault system along which the Agrio caldera, Vilú Mallín, and Domuyo volcanic complexes are aligned. Overall, the backarc in this region was affected by compression in the late Miocene and extensional collapse and transpressional deformation due to strain partitioning in the late Pliocene to Quaternary. Keywords: Upper Pliocene, Pleistocene volcanism, Southern volcanic zone, Patagonia, retroarc deformation.
*E-mail:
[email protected].
Miranda, F., Folguera, A., Leal, P.R., Naranjo, J.A., and Pesce, A., 2006, Upper Pliocene to Lower Pleistocene volcanic complexes and Upper Neogene deformation in the south-central Andes (36°30′–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. 287–298, doi: 10.1130/2006.2407(13). For permission to copy, contact
[email protected]. ©2006 Geological Society of America. All rights reserved.
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INTRODUCTION The Southern volcanic zone between 33° and 46°S occurs above a 30°E-dipping segment of the subducting Nazca plate beneath the South American plate (e.g., Bohm et al., 2002). On a broad scale, volcanic rocks have been continuously generated along this segment since the early Miocene, when the trenchnormal component of subduction increased as a consequence of the breakup of the Farallon plate (Pardo Casas and Molnar, 1987). Among these volcanic rocks are series of stratovolcanoes, calderas, and other minor volcanic centers that have erupted in Pliocene to Holocene times (e.g., Muñoz and Stern, 1988). López-Escobar et al. (1995) and Lara et al. (2001), among many others, have studied the petrology, geochemistry, and host structure of the Upper Pleistocene to Holocene volcanic rocks in the Southern volcanic zone frontal arc region. Muñoz and Stern (1988) studied Upper Pliocene to Lower Pleistocene volcanic rocks east of the modern volcanic front south of 37.5°S. In contrast, Upper Pliocene to Lower Pleistocene volcanic rocks east of the present volcanic front between 36.5°S and 38°S are just beginning to be studied in detail (e.g., Lara and Folguera, this volume, chapter 14; Kay et al., this volume, chapter 2; Varekamp, this volume, chapter 15; and references in these papers). Published and new geochronological data on a series of stratovolcanoes, volcanic complexes, and calderas aligned in a narrow longitudinal band from 36°30′S to 38°S on the eastern side of the Andes indicate the presence of an Upper Pliocene to Lower Pleistocene extinct volcanic front (Muñoz and Stern, 1988; Muñoz Bravo et al., 1989), or at least a broadened one in comparison with the 2–0 Ma volcanic arc (Lara and Folguera, this volume, chapter 14). Unlike the centers at the arc front on the western side of the Andes, which are not associated with the
Isotopic age (Ma)
Method
Rock
GEOLOGICAL SETTING AND REGIONAL GEOLOGY The area under study is located in the transition between the southern Central Andes (27°–38°S) and the northern Patagonian Andes (38°–45°S) (Fig. 1). The southern Central Andes are characterized by a recent uplift history related to the Neogene stacking of crustal thrust sheets that have partially obliterated older deformational features (Ramos, 1999). In contrast, the deformation pattern of the northern Patagonian Andes to the south shows an evolution marked by periods of foreland progression of east-vergent Late Cretaceous to Upper Miocene thrusts (Ramos, 1977; Zapata et al., 1999; Zapata and Folguera, 2006), and little Upper Neogene to Quaternary exhumation in
TABLE 1. K-Ar AGES Analytical data Location Source 40 40 K content K Ar rad. Ar atm. (%) (mol/g) (%) 0.881 n.a. 0.232 61 Vilú Mallín caldera basement SERNAGEOMIN Proyecto Riesgo nL/g Volcánico (2311) 0.944 n.a. 0.116 66 Vilú Mallín postcaldera SERNAGEOMIN Proyecto Riesgo nL/g monogenic flow Volcánico (2311) 1.098 n.a. 0.169 85 Vilú Mallín precaldera sequence SERNAGEOMIN Proyecto Riesgo nL/g Volcánico (2311) 88 Domuyo dome Miranda (1996) 2.44 7.283 × 10–8 0.107 × 10–10 mol/g n.a. n.a. n.a. n.a. Cajón Negro Fm. left side Pesce (1983, 1987) Atreuco stream n.a. n.a. n.a. n.a. Dome (Mt. Domo) Pesce (1983, 1987) n.a. n.a. n.a. n.a. Sierra de Flores Formation lava Pesce (1987) 40
6.8 ± 0.4
K-Ar
Andesite
3.1 ± 0.2
K-Ar
Andesite
4.0 ± 0.5
K-Ar
Andesite
2.5 ± 0.5
K-Ar (whole rock) K-Ar (whole rock) K-Ar (?) K-Ar (?)
Granophyre
10 ± 1 14 ± 2 0.72 ± 0.1 4.0 ± 1.0
Pliocene-Pleistocene orogenic front, these Upper Pliocene to Lower Pleistocene volcanic centers are close to the emergent fold-and-thrust belt and provide a means to constrain the age of deformation in this part of the modern retroarc. Llambías et al. (1978b), Pesce (1981, 1987), Brousse and Pesce (1982), Rovere (1993, 1998), Miranda (1996), Vattuone and Latorre (1998), and Ré et al. (2000) have presented petrological, geochemical, and geochronological studies of the volcanic rocks in these extinct retroarc centers. However, no integrated studies have been done at a regional scale with regard to: (1) the structures that host these volcanic rocks, (2) the effects of younger deformation on these centers, or (3) the temporal constraints that these volcanic rocks can place on earlier periods of deformation. The purpose of this paper is to describe and present maps of these centers, to discuss their ages and the deformation that affects them in the context of the new K-Ar ages in Table 1 and the fission-track ages in Table 2, and to put these centers into the context of the regional deformation pattern.
Andesite Dacite Andesite
Isotopic age (Ma)
Mineral
Rock
0.11 ± 0.02 0.29 ± 0.07
Zircon Zircon
Perlite Perlite
0.55 ± 0.10
Zircon
Perlite
24 (1) 23 (2) 21 23
TABLE 2. FISSION-TRACK AGES Test of fission-track data Spontaneous tracks Induced tracks Coefficient of Average Coefficient of Average Coefficient variation of variation of variation measured area 0.7 1.053 332.3 0.191 0.069 1.5 1.267 203.3 0.334 0.257 1.0 1.023 199.8 0.362 0.248 2.0 0.839 204.3 0.248 0.133
Note: Data here are reproduced from open file report (JICA, 1983).
Location Relative standard deviation 0.024 0.278 0.242 0.189
F test F value nL F (0.05) n2 1.38 2.01 2.08 2.05 1.03 2.12 1.27 2.05
Dome (Mt. Domo) Dome (Mt. Covunco)
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Figure 1. Map of part of the south-central Andes showing the late Quaternary Southern volcanic zone active front represented by the Chillán, Antuco, and Callaqui volcanic centers and the Upper Pliocene to Lower Pleistocene centers in the modern retroarc in Neuquén Province in Argentina. Note the general coincidence between the retroarc centers and the main fault systems known in the region.
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the frontal parts of the outer (eastern) part of the Neuquén foldand-thrust belt (Ramos and Barbieri, 1989). At these latitudes, a series of volcaniclastic sequences has been associated with depocenters on both sides of the Neuquén Andes (Fig. 2). The youngest depocenter, entirely on the western side of the Andes and with limited exposures along the axial part of the cordillera, was generated between 15 and 10 Ma in the southern Cura Mallín basin (Suárez and Emparán, 1995, 1997). Rocks with similar ages occur on the eastern side of the Andes. Among these are volcaniclastic sequences in the Cajón Negro Formation (Pesce, 1981, 1987), which are associated with a depocenter bounded on the south by the Nahueve fault (Fig. 1). Radiometric ages indicate an age range from 14 to 10 Ma
for the Cajón Negro Formation (Pesce, 1987). This range could be extended to 9 Ma based on an age reported by Burns (2002) east of the study area (Table 1). An equivalent unit, the Charilehue Formation (Uliana et al., 1973) extends into the Chos Malal fold-and-thrust belt (Fig. 1). Younger volcanic units, represented by the Cola de Zorro (Niemeyer and Muñoz, 1983) and Malleco Formations (Suárez and Emparán, 1997), are widely distributed along the axial part of the Andes. Their ages are between 6 and 3.5 Ma (Niemeyer and Muñoz, 1983; Muñoz and Stern, 1988; Muñoz Bravo et al., 1989; Linares et al., 1999). These volcanic rocks appear to have largely erupted from fissures, as only a few volcanic centers of this age have been identified (see Lara and Folguera, this vol-
Figure 2. Stratigraphic chart showing the Neogene volcanic units in the arc and the retroarc that are discussed in the text.
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Upper Pliocene to Lower Pleistocene volcanic complexes ume, chapter 14). One of these centers is the prominent Sierra de Flores volcanic center (Pesce, 1987) in the study area (Fig. 1). The youngest centers in the region are the Upper Pliocene to Quaternary volcanic complexes described herein, which are distributed along the eastern slope of the Andes between 36°30′S and 38° S (Fig. 1). From north to south, the main centers are the Domuyo volcanic complex, the Los Cardos–Centinela stratovolcano, the Trohunco caldera, the Vilú Mallín caldera, and the Agrio caldera. UPPER PLIOCENE TO LOWER PLEISTOCENE VOLCANIC CENTERS Cerro Domuyo Area The Cerro Domuyo area (36°38′S, 70°26′W) includes one of the most important igneous centers of the northern Neuquén Andes (Fig. 1). This area has been studied from a regional geologic point of view by Groeber (1947), Llambías et al. (1978a, 1978b), Pesce (1981), and Brousse and Pesce (1982), and from a geothermal point of view by Jurio (1978), Palacios and Llambías (1978), JICA (1983), Pesce (1983, 1987), and Panarello et al. (1990).
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Cerro Domuyo (4709 m) (Fig. 3) itself is a dome that is unconformably emplaced in an anticline. The anticline (Groeber, 1947) folds Permian-Triassic deposits of the Choiyoi Group and Upper Triassic to Upper Cretaceous sedimentary rocks of the Neuquén Basin (Fig. 2). The dome in the summit area of Cerro Domuyo is composed of high-K rhyolite with a porphyric and granophyric texture (Miranda, 1996). Llambías et al. (1978a, 1978b) postulated a late Miocene age for this dome based on stratigraphic considerations. The new K-Ar age of 2.5 ± 0.5 Ma in Table 1 shows that the Domuyo dome is actually late Pliocene in age (Miranda, 1996). Other Cenozoic rocks in the Cerro Domuyo area are mainly volcanic in origin (Llambías et al., 1978b; Brousse and Pesce, 1982; Pesce, 1987). The oldest sequences are in the Charilehue Formation (Fig. 2), which is composed of basaltic andesitic to andesitic flows (Uliana et al., 1973; Llambías et al., 1978b). These flows are folded into an anticline enclosing the Domuyo intrusive dome. Everywhere, the Charilehue volcanic rocks rest in angular unconformity on folded Mesozoic sedimentary rocks (Llambías et al., 1978b). Pesce (1981) correlated the Charilehue Formation with the Cajón Negro Formation to the west (Fig. 1), the age of which is constrained by K-Ar ages of 14 ± 2 Ma and 10 ± 1 Ma. The Cajón Negro Formation
Figure 3. Map of the Domuyo volcanic complex and surrounding region showing the distribution of Permian-Triassic to Pleistocene volcanic rocks and the principal structures in the region. The arrow points to the location of the new K-Ar age in Table 1.
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extends to the international border where the sequence is gently folded. The age of the Cajón Negro Formation is further constrained by the overlying andesitic lava flows of the middle to late Miocene Quebrada Honda Formation (Pesce, 1981) and the 40Ar/ 39Ar age of 9 Ma on a flow in the Nahueve valley (Burns, 2002; Burns et al., this volume, chapter 8). Horizontally and unconformably overlaying the Charilehue Formation are the lavas flows of the Sierra de Flores Formation (Brousse and Pesce, 1982). These 1.5–2-m-thick basaltic andesitic and andesitic flows were erupted from a volcanic center west of Cerro Domuyo (Fig. 1). Their age is constrained by a K-Ar age of 4 ± 1.0 Ma (Pesce, 1987). Late Pliocene to Pleistocene volcanic activity in the Domuyo region is characterized by the phreatomagmatic rhyolitic deposits and extrusive domes that occur southwest and on the northeastern flanks of the Domuyo summit (Fig. 3). These domes, which were called the Magmatismo Dómico by Brousse and Pesce (1982), were emplaced along NE– and E-W–oriented extensional faults (Pesce, 1987). Their ages are constrained by K-Ar ages, which range from 720 to 110 ka (Table 1), and Pleistocene fission-track analyses on zircons, which are listed in Table 2 (Brousse and Pesce, 1982; JICA, 1983). Llambías et al. (1978b) inferred two stages of development for the Domuyo anticline. The first involved folding of the sedimentary rocks and was considered to have occurred near the end of the Cretaceous. A second, milder phase of deformation was inferred to follow erosion of the Cretaceous sequences, postdate deposition of the Miocene Charilehue Formation, and to predate the intrusion of the Domuyo dome. There is still some uncertainty as to whether the second deformation phase is solely related to the viscous emplacement of the Domuyo dome. If the folding is regional, the deformation fits with a pulse of Neogene contractional deformation during the 14–10 Ma age assigned to the Charilehue lavas and eruption of the Sierra de Flores flows at 4 ± 1.0 Ma. Support for regional deformation near this time comes from a 12 Ma fission-track age for uplift in Burns (2002) on the eastern slope of the Cordillera del Viento, immediately south of the Domuyo area. Other support comes from the argument of Kay et al. (this volume, chapter 2) that compressional deformation occurred between 11 ± 0.2 Ma and 4.0 ± 0.4 Ma just east of the Cordillera del Viento. A further constraint on the youngest age of deformation comes from the Pleistocene domes of the Domuyo complex. Los Cardos–Centinela Volcanic Center To the south of the Domuyo center is the Los Cardos– Centinela volcanic center (37°06′S, 70°52′W) (Figs. 1 and 4). The main edifice is a stratovolcano that erupted olivine and plagioclase-bearing basalts and subordinate pyroclastic deposits. The age of this center has been constrained between 3.2 and 2.5 Ma (Rovere, 1993, 1998). The eastern face of the stratovolcano shows minimal erosion, whereas the western slope has been heavily affected by Holocene mass wasting (González Díaz et al., 2005). The top of the center contains a summit caldera that has a poorly constrained age. Postdating the caldera is a series of minor
preglacial stratovolcanoes that erupted near the apex. Moderate amounts of postcaldera dome activity also occurred on the eastern flank (Fig. 4). Based on the pillow-like structures in the youngest flows, the latest volcanic activity is considered to have occurred during synglacial times. In analogy with the glacial history of the Chillán volcano (Dixon et al., 1999) to the west (Fig. 1), the latest volcanic activity would have occurred after 30 ka. A Miocene age for the principal deformation in this region can be inferred from the fact that the 3.2–2.5 Ma lavas of the Los Cardos–Centinela stratovolcano, along with underlying early Pliocene Cola de Zorro volcanic rocks, lie in angular unconformity over deformed strata of the Miocene Cura Mallín Formation. The youngest age of deformation is constrained by deformed avalanche deposits west of the center. The age of these deposits is considered to be younger than 30 ka, based on the absence of glacial erosive features on the avalanche deposits. Folding and reverse faulting in these deposits show a N to NW trend (Folguera et al., this volume, chapter 11). Trohunco Caldera Area Farther south is the partially eroded Trohunco caldera (37°18′S, 71°01′W), which is located on the eastern side of the Andes, west of the Loncopué trough (Fig. 1). Volcanic rocks from this center are largely andesitic breccias that have porphyritic textures and contain vesicles up to 8 mm in diameter. The phenocrysts, which make up 45% of the rock, are mainly plagioclase (40%), augite (5%), and accessory small opaque minerals. This 15-km-diameter caldera is comparable to the similarsized Agrio caldera farther south (Fig. 1) in that its rim (precaldera units) is formed by Lower Pliocene volcanic rocks of the Cola de Zorro Formation. It differs in not having a resurgent facies. However, these facies could have been removed, since the eastern half of the caldera has been eroded at the orogenic front (Fig. 5). Volcanic rocks outside of the caldera have yielded K-Ar ages of 3.6 ± 0.2 and 3.6 ± 0.5 Ma (Muñoz Bravo et al., 1989). The relationship between this volcanic center and surrounding volcanic units shows that two main pulses of contractional deformation occurred in the area. The oldest one is indicated by an angular unconformity that separates folded sequences of early Miocene age from subhorizontal lava flows dated at 3.6 Ma. A younger contractional deformation is indicated by Pliocene intracaldera volcanic rocks near the eastern edge of the caldera that are affected by gentle folding. These rocks are in turn overridden by Lower Miocene volcanic rocks in a NW-oriented high-angle reverse fault (Fig. 5). This deformation appears to be associated with a system of N-NW–trending faults and folds that Folguera et al. (this volume, chapter 11) consider to be Quaternary in age Vilú Mallín Caldera Area The 6–7-km-diameter Vilú Mallín caldera (37°28′S, 70°45′W) (Fig. 6) is located in the northern part of the Quaternary Loncopué trough to the east of the Trohunco, Los Cardos–
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Upper Pliocene to Lower Pleistocene volcanic complexes
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Figure 4. Map of the Pliocene to Pleistocene Centinela volcanic complex showing the distribution of volcanic and subvolcanic units relative to reverse faults and caldera collapse.
Centinela, and Agrio calderas (Fig. 1). The principal volcanic rocks are basaltic andesites, which have plagioclase and augite phenocrysts that compose ~20% of the rock. The caldera was formed by the collapse of a plateau made up of a precaldera sequence that has yielded a K-Ar age of 4 ± 0.5 Ma (Table 1). The southern part of the depression has been obliterated by postcaldera monogenetic pulses of basic lavas, one of which yielded a K-Ar age of 3.1 ± 0.2 Ma (Table 1). Resurgent activity has also formed a series of basaltic domes that erupted from annular rings around the edge of the caldera. The Vilú Mallín volcanic center, the Mandolegüe volcanic field, and Trocomán volcano, are aligned in a NE-trending volcanic chain that is controlled by the Trocomán dextraltranstensional fault (Figs. 1 and 6). An apparent displacement in the Vilú Mallín caldera rim (Fig. 6) suggests probable activity on this fault during the last 4 m.y. The volcanic basement of these three complexes is an andesitic plateau, which has yielded a K-Ar age of 6.8 ± 0.4 Ma (Table 1). This volcanic basement horizontally covers folded strata of Lower Miocene age. Observations from the Vilú Mallín center and the underlying 6.8 ± 0.4 Ma andesitic plateau help to constrain the age of Neogene deformation in the area. A first observation is that the
basal lavas of the Vilú Mallín center erupted over a regional angular unconformity on sedimentary beds of the Lower Miocene Cura Mallín Formation in the Reñileuvú valley (Fig. 6). A second is that the Vilú Mallín center is cut by the Trocomán valley (Fig. 6), which is the morphological expression of the NE-trending Trocomán fault. This fault system includes a series of along-strike pull-apart basins that formed in response to dextral displacement (Folguera et al., 2004). The basal volcanic rocks of the Vilú Mallín caldera occur in a small pull-apart basin formed directly along the Trocomán fault trace (Figs. 1 and 6). Despite evidence for strike-slip motion, the main effect of the Trocomán fault is the extensional faulting that led to the collapse of the southern side of the caldera (Fig. 6). The Agrio Caldera The Agrio caldera (37°51′S, 70°26′W) (Figs. 1 and 7) is a quadrangular (15 × 20 km) depression filled by volcaniclastic successions with ages ranging from 2.5 Ma to younger than 30 ka (Pesce, 1989; Linares et al., 1999; Melnick and Folguera, 2001). The volcanic rocks vary from andesitic to basaltic in composition. Most have porphyritic texture and typically contain 25% plagioclase and 5% augite phenocrysts. Aphanitic textures are also present.
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Figure 5. Map of the Pliocene to Pleistocene Trohunco caldera showing the crater relative to the distribution of Miocene and Pliocene-Quaternary volcanic rocks and avalanche deposits in the region and the reverse faults that cut them.
Figure 6. Map showing the distribution of Miocene and Pliocene-Quaternary volcanic units in the region of the Vilú Mallín caldera and Trocomán volcanic complex and the extensional faults that cut them. The arrows point to the locations of the new K-Ar ages in Table 1.
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Upper Pliocene to Lower Pleistocene volcanic complexes The collapse that created the Agrio caldera took place in more than one episode (Melnick and Folguera, 2001). The oldest syncaldera volcanic rocks that are associated with the earliest collapse are part of the Las Mellizas Formation (Pesce, 1989). The Las Mellizas Formation, which has been dated at 2.5 Ma (Pesce, 1989), covers most of the caldera. Its main depocenter is near the northwestern half of the volcanic depression (Fig. 7). The age of the second collapse is bracketed between 1.6 and 0.8 Ma (Pesce, 1989). This collapse is restricted to the northern part of the caldera, where it occurred along a W- to NW-trending extensional fault system (Folguera and Ramos, 2000). Similarly oriented extensional fault systems in the southern part of the caldera have controlled the emplacement of volcanic rocks with ages from 1.2 Ma to synglacial. The 1.2 Ma age corresponds to the basal lavas of the Copahue volcano, which fill two extensional depocenters. One center is in the western part of El Agrio graben, and the other is a small pull-apart basin along the upper Lomín River (Melnick and Folguera, 2001). Synglacial volcanic rocks, the ages of which have been inferred from pillowlike structures related to flow under the ice, are systematically controlled by W- to NW-trending structures (Melnick et al., this volume, chapter 4). Postglacial volcanic rocks in the Copahue volcano and to the north (Fig. 7) are controlled by NE-trending extensional faults (Folguera and Ramos, 2000). The Agrio caldera shows the relations between several pulses of deformation, which can also be seen in neighboring volcanic centers. The youngest precaldera sequences in the area have ages of 5–4 Ma (Linares et al., 1999) and belong to the Cola de Zorro Formation (Niemeyer and Muñoz, 1983). These rocks unconformably cover the Lower Miocene Cura Mallín
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Formation, indicating a widespread pulse of post–early Miocene compressive deformation. The Agrio caldera is a transtensional basin that is controlled by W-NW– and NE-trending normal faults (Fig. 7). These faults form an almost rhombohedral depocenter that is internally segmented by W-NW–trending faults. Well-dated synextensional volcanic rocks indicate recurrent collapse in the area. The caldera and other transtensional depocenters in the region were active in the late Pliocene after the Cola de Zorro eruptions, and continued to be active through the Quaternary until postglacial times. The main regional structures controlling the collapse of the Agrio caldera area can be seen in Figure 1. These are: (1) the northernmost end of the Liquiñe-Ofqui fault zone, which runs through the Upper Pleistocene to Holocene volcanic front to the south (Lavenu and Cembrano, 1999), and (2) the southern end of the Antiñir-Copahue fault system, which runs through the inner retroarc zone (Folguera et al., 2004). Both fault systems have a dextral component associated with differing amounts of extension and compression as the faults change orientation. The clockwise step design of these two fault-system traces is compatible with the development of transtension in the Agrio caldera area (Fig. 7). DISCUSSION AND CONCLUSIONS Fission-track ages of 12 Ma (Burns, 2002) from the Cordillera del Viento, immediately to the south of Cerro Domuyo (Fig. 1), show that the youngest major uplift in the region occurred in the late Miocene. Based on geochronological data and field relationships between the Domuyo igneous complex and its basement, two pulses of deformation can be identified in
Figure 7. Map of the El Agrio caldera region showing the distribution of Miocene and Pliocene-Quaternary volcanic rocks relative to avalanche deposits and the major structures in the region. See text for discussion.
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Figure 8. Regional map showing the interplay of Pliocene-Quaternary extensional and contractional deformational systems in the retroarc. The Los Cardos–Centinela and Trohunco centers occur in the region where contraction has occurred along N-NW–oriented faults. Cerro Domuyo, Vilú Mallín, and the Agrio caldera occur in the region where extension has occurred along NE-trending fault systems. See text for further discussion.
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Upper Pliocene to Lower Pleistocene volcanic complexes the last 15 m.y. that appear to be of regional extent. The first produced the folding in the Charilehue Formation after 10–9 Ma, and the second, the extension and the collapse of the northern part of the Cordillera del Viento after 1 Ma. The first pulse that folded the Charilehue Formation must be younger than 10–9 Ma if the age correlation with the Cajón Negro Formation is correct. This pulse is either younger than the 12 Ma uplift proposed by Burns (2002), or the fissiontrack age does not accurately reflect the last uplift, or the folding of the Charilehue Formation was not associated with substantial uplift. Based on field relations with the 4 Ma Sierra de Flores Formation lavas that horizontally cover and unconformable overlie middle Miocene lava flows, the pulse that folded the Charilehue Formation can be constrained between 9 and 4 Ma. The analysis of the Vilú Mallín and Agrio calderas (Figs. 6, 7, and 8) puts other constraints on the age of contractional deformation in the region. Constraints on the upper age of deformation at these centers are 6.8 Ma and 5 Ma, respectively. Taken together with the constraints in the Domuyo region, regional-scale contraction likely occurred between 9 and 6.8 Ma. The second pulse of deformation involved extension and collapse. This pulse is shown in the Domuyo area by the emplacement of the Pleistocene domes that were favored by extensional structures. Extensional deformation is also recorded in the Vilú Mallín and Agrio caldera areas, where NE-trending extensional fault systems were active during late Pliocene to Quaternary times (Fig. 8). In contrast, the Los Cardos–Centinela volcanic complex and Trohunco caldera show a different style for the youngest deformational pulse. In the Los Cardos–Centinela center, N-NW– trending structures indicate contraction during the late Quaternary. At the Trohunco caldera, the N-NW–trending system shows evidence for contractional deformation in the late Pliocene to Quaternary. Based on these differences, the volcanic centers can be put in two groups. The first includes the Agrio and Vilú Mallín calderas and the Domuyo volcanic complex, and the second the Centinela and Trohunco centers. The first group erupted in relation to NE-trending Pliocene to Quaternary extensional fault systems (Figs. 6, 7, and 8), and the second is associated with NW-trending contractional faults active in Quaternary times. On a regional scale, this difference reflects the inhomogeneous nature of young deformation in the area and the strain partitioning between faults, which accommodates extension and contraction at the same time. ACKNOWLEDGMENTS The authors would like to thank Adriana Bermúdez and Eduardo Llambías for their reviews of an earlier version of this manuscript. This study was made possible by funding from PICT 06729/99 of Agencia Nacional de Promoción Científica y Tecnológica to V.A. Ramos. We thank Suzanne Mahlburg Kay for several reviews and suggestions that substantially improved the presentation and clarified the concepts in this paper.
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Pesce, A.H., 1983, Evaluación geotérmica del área del cerro Domuyo (Neuquén): SEGEMAR (Servicio Geológico Minero Argentino) Argentina Open-File Report, Serie Contribuciones Técnicas, Geotermia, v. 1, 46 p. Pesce, A.H., 1987, Evaluación geotermica del area Cerro Domuyo, Provincia del Neuquén, República Argentina: Revista Brasileira de Geofísica, v. 5, p. 283–299. Pesce, A., 1989, Evolución volcano-tectónica del complejo efusivo CopahueCaviahue y su modelo geotérmico preliminar: Revista de la Asociación Geológica Argentina, v. 44, no. 1–4, p. 307–327. Ramos, V.A., 1977, Estructura de la Provincia de Neuquén, in Rolleri, E.O.. eds., Geología y recursos naturales de la Provincia del Neuquén: 7th Congreso Geológico Argentino (Neuquén): Buenos Aires, Asociación Geológica Argentina, p. 9–24. Ramos, V.A., 1999, Plate tectonic setting of the Andean Cordillera: Episodes, v. 22, no. 3, p. 183–190. Ramos, V.A., and Barbieri, M., 1989, El volcanismo Cenozoico de Huantraico: Edad y relaciones isotópicas iniciales, Provincia del Neuquén: Revista de la Asociación Geológica Argentina, v. 43, p. 210–223. Ré, G.H., Geuna, S.E., and López Martínez, M., 2000, Geoquímica y geocronología de los basaltos de la región de Aluminé (Neuquen-Argentina), in Proceedings, 9th Congreso Geológico Chileno: Puerto Varas, Servicio Nacional de Geología y Minería de Chile, v. 2, no. 6, p. 62–66. Rovere, E., 1993, K/Ar ages of magmatic rocks and geochemical variations of volcanics from South Andes (37° to 37°15′S–71°W), in Proceedings, 2nd Japan Volcanological Society: p. 107. Rovere, E., 1998, Volcanismo Jurásico, Paleógeno y Neógeno en el noroeste del Neuquén, Argentina, in Proceedings, 10th Congreso Latinoamericano de Geología: Buenos Aires, Servicio Geológico Minero Argentino, v. 1, p. 144–149. Suárez, M., and Emparán, 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. Suárez, M., and Emparán, C., 1997, Hoja Curacautín. Regiones de la Araucanía y del Bío Bío: Carta Geológica de Chile: Santiago, Servicio Nacional de Geología y Minería de Chile, v. 71, p. 105, scale 1:250,000, 1 sheet. Uliana, M., Dellape, D., and Pando, G., 1973, Estratigrafía, estructura y posibilidades petroleras del extremo noroeste de la Provincia de Neuquén: Buenos Aires, Yacimientos Petrolíferos Fiscales Open-File Report. Varekamp, J.C., Maarten deMoor, J., Merrill, M.D., Colvin, A.S., Goss, A.R., Vroon, P.Z., and Hilton, D.R., 2006, this volume, Geochemistry and isotopic characteristics of the Caviahue-Copahue volcanic complex, Province of Neuquén, Argentina, 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, doi: 10.1130/2006.2407(15). Vattuone, M., and Latorre, C., 1998, Caracterización geoquímica y edad K/Ar de basaltos del Terciario superior de Aluminé, Neuquén, in Proceedings, 5th Congreso Latinoamericano de Geología: Buenos Aires, Servicio Geológico Minero Argentino, v. 2, p. 184–190. Zapata, T., and Folguera, A., 2005, Tectonic evolution of the Andean fold and thrust belt of the southern Neuquén Basin, Argentina, in Veiga, G.D., Spalletti, L., Howell, J.A., and Schwarz, E., eds., The Neuquén Basin: A case study in sequence stratigraphy and basin dynamics: Geological Society of London Special Publication 252, p. 37–56. Zapata, T., Brissón, I., and Dzelalija, F., 1999, The role of basement in the Andean fold and thrust belt of the Neuquén Basin, in Proceedings, Thrust Tectonics (third): London, University of London, p. 122–124.
MANUSCRIPT ACCEPTED BY THE SOCIETY 22 DECEMBER 2005
Printed in the USA
Geological Society of America Special Paper 407 2006
The Pliocene to Quaternary narrowing of the Southern Andean volcanic arc between 37° and 41°S latitude Luis E. Lara* Servicio Nacional de Geología y Minería (SERNAGEOMIN, Chile), Ave. Santa María 0104, Providencia, Santiago, Chile, and IRD-LMTG, Laboratoire des Mécanismes de Transfert en Géologie, Université Paul Sabatier-Toulouse III, 14 avenue Edouard Belin, 31400, Toulouse, France Andrés Folguera* Laboratorio de Tectónica Andina, Universidad de Buenos Aires, Consejo Nacional de Investigaciones Científicas y Técnicas, Buenos Aires, Argentina ABSTRACT A complex arc-backarc system developed at the western margin of the Neuquén Basin north of 37°S during the late Cenozoic. Both transpressional deformation on the western orogenic front (Liquiñe-Ofqui fault system) and contractional deformation of the eastern Andean foothills occurred during the middle to late Miocene. South of 38°S, a physiography dominated by uplifted blocks and elongated basins was the site of intense volcanism from the early Pliocene to the Holocene. A wide volcanic arc was established from the western orogenic front to the eastern foothills of the Andes during the Pliocene. This volcanic phase was coeval with both extended transpressional deformation along the frontal arc and transient extensional episodes in the inner retroarc. Arc-front geochemical signatures of the magmas occurred further east of the front showing an increased subduction input in the subarc mantle. A decrease in plate convergence velocity 2–3 m.y. ago, along with a stable arc front caused a progressive westward narrowing of the Quaternary volcanic arc, probably since ca. 1.6 Ma. From the middle to late Pleistocene, volcanism was mainly centered around the Liquiñe-Ofqui fault system with minor Holocene activity in the eastern Andean region. Arc-front geochemical signatures are now restricted to the present volcanic front. Morphologically, this has resulted in paired volcanic belts that reflect different stages of arc narrowing rather than separate arc fronts. Keywords: Pliocene, Quaternary volcanic arc, arc narrowing, Southern Andes. INTRODUCTION The Mesozoic evolution of the Neuquén Basin is widely understood due to the abundant literature that has been stimulated by the oil industry (Vergani et al., 1995; Kozlowski et al., 1996; *E-mails:
[email protected];
[email protected].
Zapata et al., 1999, 2002). For similar reasons, the Tertiary histories of the intra-arc basins and alluvial systems developed on its southwestern margin are also well known (Jordan et al., 2001; Burns, 2002; Burns et al., this volume, chapter 8). However, the tectonic response of the western margin of the Neuquén Basin to Andean orogenesis and the ongoing volcanism has been studied only in recent times, and integrated models for Cenozoic arc and
Lara, L.E., and Folguera, A., 2006, The Pliocene to Quaternary narrowing of the Southern Andean volcanic arc between 37° and 41°S latitude, 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. 299–315, doi: 10.1130/2006.2407(14). For permission to copy, contact
[email protected]. ©2006 Geological Society of America. All rights reserved.
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foreland deformation are just emerging (e.g., Kay et al., this volume, chapter 2). The evolution of the volcanic arc on the western and eastern sides of the Andean range after the main phase of Andean uplift between ca. 19 and 5 Ma is not well known, and models for the Pliocene to Quaternary magmatic arc in an acrossorogen perspective are needed. North of 38°S, migration of the volcanic arc, along with steepening of the slab have been proposed as explanations for the changing physiognomy and geochemical signature of Pliocene to Quaternary Andean magmas (Stern, 1989; Muñoz and Stern, 1988; Kay et al., 2005). South of 38°S, published structural and chemical data, along with new field information, suggest a changing architecture for the volcanic arc in the late Pliocene or early Pleistocene. This situation provides an opportunity for testing the relation between convergence rates, the tectonic regime of the upper crust, and the architecture of volcanic arcs in a continental margin. In this paper, we analyze the spatial and temporal evolution of the Pliocene and Quaternary volcanic arc on the southwestern margin of the Neuquén Basin, south of 38°S. To address the architecture of a volcanic arc and its evolution, it is necessary to constrain some margin-scale features of the subduction zone and the behavior of the magma source region. The first-order distribution of volcanic centers along a volcanic arc is related to the subduction angle and the distance to the trench because of the pressure-dependent locus of the melting zone. The tectonic regime of the upper crust controls the ascent and local emplacement of magma batches, their differentiation processes, and their eruptive style. As observed by Marsh (1979), a pair of volcanic chains within a single volcanic arc is a common feature that can be recognized in arcs like the Aleutians, Kamchatka, Kurile, NE Japan, Indonesia, and Scotia. Two dehydration fronts above the slab have been proposed as the cause of paired volcanic belts (Tatsumi and Eggins, 1995). In most cases, the depth to the top of the subducted slab beneath the trench-side volcanic chain is ~100 km, and it is ~200 km below the backarc-side chain (Tatsumi and Eggins, 1995). The amount of volatiles and pressure-temperature (P-T) conditions of the mantle wedge influence the melting degree, which is higher on the trench side than the backarc side. Here, we use field evidence, geochemistry, and published K-Ar and new 40Ar/ 39Ar data to present a temporal overview of the evolution of Pliocene to Quaternary volcanic arcs on the southwestern margin of the Neuquén Basin. MORPHOSTRUCTURAL FEATURES AND TECTONIC EVOLUTION Between 38°S and 41°S, the South American margin shows several first-order features: (1) a partially filled trench with an active accretionary prism; (2) forearc basins with Cretaceous to Quaternary infill on the continental shelf; (3) a coastal range formed by Paleozoic-Triassic rocks, which show evidence of polybaric metamorphism as young as Jurassic (e.g., Duhart et al., 2001); (4) a central valley, the infill of which
includes Tertiary sedimentary and volcanic sequences covered by Quaternary glacial and pyroclastic deposits; (5) a ridge (Loncoche Ridge; Chotin, 1975) that cuts the central valley from the coastal range to the Andean western front; (6) the main Andean cordillera, where the active volcanic front is located and the basement consists of Mesozoic-Cenozoic plutons of the North Patagonian Batholith, roof pendants of Mesozoic to Tertiary volcanic and sedimentary sequences, and a Quaternary volcanic and glacial cover; (7) an inner retroarc region at the eastern foothills of the main cordillera, where uplifted Mesozoic blocks are separated by flat valleys with Pliocene to Quaternary volcanic infill; and (8) a west-verging fold-and-thrust belt developed throughout the retroarc area. Main Cordillera The Main Cordillera of the Southern Andes forms a narrow, yet imposing mountain belt that reaches a mean altitude of ~2500 m at 38°S and decreases in elevation southward. Marginscale faults cut and parallel the main Andean orogen (Fig. 1) at this latitude. The most important is the northern segment of the Liquiñe-Ofqui fault system (Lavenu and Cembrano, 1999a; Cembrano et al., 2002), which trends N-NE along the Main Cordillera to 38°S, where it intersects NW-trending faults (Fig. 1). This fault system has been active since the late Miocene, but could have had activity as early as the Cretaceous (Cembrano et al., 2000, 2002). In addition, oblique northwest pre–late Miocene structures limit well-defined geologic domains (Fig. 1). These pre-Andean structures were reactivated during both a Mesozoic episode of crustal attenuation and in the OligoceneMiocene development of intra-arc basins during a period of fast convergence (Jordan et al., 2001). Quaternary reactivation of these oblique structures is suggested by the seismic pattern in the coastal area (Bohm et al., 2002) and historical earthquaketriggered eruptions in the volcanic arc (Lara et al., 2004a). Inner Retroarc—The Copahue–Pino Hachado Block The eastern Andean foothills have a basin-and-range–type topography, as shown by the well-defined pre-Andean Copahue– Pino Hachado block along the drainage divide (Fig. 1). This west-verging uplifted block is related to the inversion of a late Oligocene half-graben. Oligocene to early Miocene volcanic rocks of the Cura Mallín Formation (Suárez and Emparán, 1997), which erupted in association with the half-graben, are gently folded along the eastern and highest flank of the block near the Pino Hachado caldera. Jurassic basement is exposed on the western flank. Graben inversion must have taken place before ca. 1.8 Ma, because the basal lavas of the Pino Hachado caldera (Muñoz and Stern, 1985) seal the structure. Inversion probably occurred before the deposition of the horizontal Meseta del Arco pyroclastic flows at ca. 4.5 Ma. Subhorizontal lava flows near the Pino Hachado pass yield older K-Ar ages (4.8 ± 0.2 Ma; Linares and González, 1990).
Pliocene to Quaternary narrowing of the Southern Andean volcanic arc
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Figure 1. Pliocene, Pliocene-Quaternary, and Quaternary volcanic belts in the Andes between 38° and 41°S at western margin of the Neuquén Basin (modified mainly from Delpino and Deza, 1995; Muñoz and Stern, 1989; Suárez and Emparán, 1997; Campos et al., 1998; Rodríguez et al., 1999; and Lara and Moreno, 2004, among others). Vents are labeled where recognizable. Main fault systems and Tertiary basins are also shown. Key: FM—Formación Malleco; BO—Bonete; TE—Trocolan; TM—Las Monjas; CT—Cerro Trolón; CP—Copahue; R—Rahue; B— Butahuao; PS—Pino Solo; PH—Pino Hachado; PM—Palao Mahuida; QM—Queli Mahuida; CB—Nevados de Caburgua; TT—Cerro Trautrén; QQ—Quinquilil or Colmillo del Diablo; LP—Laguna Los Patos; P—Paimún; C—Carirriñe; QA—Sierra de Quinchilca; HQ—Huanquihué; PO—Pirihueico; QO—Quelguenco; CH—Chihuío; CN—Cordillera Nevada; M—Mencheca; F—Fiuchá; CA—Cordón de Alvarez; MR— Mirador; PJ—Pantoja; S—Sarnoso; PD—La Picada; CO—Chapuco; HH—Hueñu-Hueñu; GD—Garganta del Diablo; CD—Cuernos del Diablo; R—Reloncaví.
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The eastern side of the Copahue–Pino Hachado block is affected by extensional structures that dip to the east, suggesting local collapse during the late Miocene to Quaternary. A lower age constraint for local extension comes from the ca. 2.3 Ma K-Ar age of the basal lavas inside the Loncopué trough (Ramos, 1977; Linares and González, 1990; Leanza et al., 2001). To the west of the Copahue–Pino Hachado uplift, the Bío Bío–Aluminé trough is bordered by two possible extensional structures that almost reach the Huincul Arch, the best-known structure in the Neuquén Basin (e.g., Mosquera and Ramos, this volume, chapter 5). Farther south, the Collón Curá and Ñirihuao basins are also northwest elongated depressions. PLIOCENE-QUATERNARY VOLCANISM The main focus of this paper is the eroded late Pliocene to middle Pleistocene stratovolcanoes that occur both along the Main Cordillera and in the eastern Andean foothills. These volcanic rocks show a wide spectrum of eruption styles, degree of erosion, and geochemical signatures. Their ages are summarized in Tables 1–3. Early Pliocene: Fissure and Shield Volcanism Early Pliocene volcanic rocks related to heavily eroded volcanic centers commonly form the base of the PlioceneQuaternary stratovolcanoes on both the western and eastern side of the Andean range. Near 38°S, a prominent 500-m-thick volcanic and sedimentary sequence (Malleco Formation of Suárez and Emparán, 1997) has a base composed of volcanic breccias, tuffs with interbedded gravels, and lavas. These predominantly basaltic to low-silica andesitic volcanic rocks have yielded K-Ar ages that range from ca. 4.4 to 2.3 Ma. The basal units are conformably overlain by younger lavas and cut by feeder dikes and remnants of vent facies. Volcanic vents have not been recognized, but preliminary studies related to polarities of local facies indicate flow direction to the west for the lower and middle members of the Malleco Formation. To the south near 40°S, a 150–500-m-thick sequence composed of basaltic lavas, breccias, and coarse gravels (Estratos de Lago Ranco of Campos et al., 1998) has been dated at ca. 5.8–2.4 Ma (Campos et al., 1998; Lara and Moreno, 2004). East of the Andean range in the Lonquimay area, subhorizontal piles of mainly basaltic lavas (Llanquén-Ranquil and Tuetué sequences; Suárez and Emparán, 1997) have Pliocene K-Ar ages from ca. 5.2 to 3.2 Ma. All of these mainly basaltic to andesitic, subhorizontal sequences have morphological features typical of effusive volcanic rocks not related to compound volcanic structures. The absence of internal unconformities within thick sequences also supports a fissure style of volcanism. Similar features occur in the thick volcanic piles north of 38°S where González and Vergara (1962) defined the Cola de Zorro Formation. South of 40°S, early Pliocene magmatic rocks in the Main Cordillera are granitoids and synplutonic mylonites related to
the North Patagonian Batholith. The absence of volcanic rocks can be attributed to high exhumation rates along the axial Andean zone as shown by apatite fission tracks and thermal histories (Gräfe et al., 2002; Cembrano et al., 2000). Late Pliocene to Early Pleistocene: Shield Volcanoes in a Broad Volcanic Arc Late Pliocene to early Pleistocene volcanic rocks partially overlie early Pliocene volcanic sequences from the main Andean range to the eastern uplifted blocks. Near the modern arc front, these volcanic rocks occur in lava flows, volcaniclastic sequences, and as deeply eroded stratovolcanoes (Figs. 1 and 2). Remnant subhorizontal or gently dipping thin flows that are mostly basaltic in composition have effusive volcanic features. The eastern belt is formed by partially preserved volcanic structures. The upper member of the Malleco Formation (38°S), which can be up to ~500 m thick, consists of basaltic andesitic lavas associated with poorly preserved necks at El Peñón (ca. 1.8– 1.3 Ma), Paso Marcial (younger than 1.8 Ma), Piedra Marcada (ca. 1.8 Ma), and Las Mellizas and Cerros de Lanco volcanoes (Suárez and Emparán, 1997). Near 39°S, several small volcanic piles that have poorly preserved vent facies cover granitoids of the North Patagonian Batholith at Cerro Trautrén (ca. 0.8 Ma), Cerro Maichin (ca. 0.9 Ma), and Laguna Los Patos and Carirriñe (Lara et al., 2001). Farther south, a thick series (~550 m) of basalts and laharic breccias composes the Estratos de Chapuco sequence (ca. 1.0–0.4 Ma) at the base of the Quaternary Osorno volcano (Moreno et al., 1985; Lara et al., 2001). The volcanic rocks in the Estratos de Hueñu-Hueñu (ca. 1.43 Ma) at the base of the Calbuco volcano (41.3°S) can be related to remnant vents (Moreno et al., 1985). The thick volcanic Garganta del Diablo sequence (Mella et al., 2005) at the base of Tronador volcano has a K-Ar age of ca. 1.3 Ma. This age could be the same as that of the Steffen volcanic complex. Better-preserved central volcanoes are also part of this group. Nevados de Caburgua (39°S) is a ring structure that has pyroclastic beds and lavas surrounding an andesitic laccolith (K-Ar ages from 2.4 to 0.8 Ma; 40Ar/ 39Ar age of ca. 984 ka). Huanquihué at 39.8°S, Pirihueico at 39.9°S (K-Ar age of ca. 1.5 Ma; 40Ar/ 39Ar age of ca. 601 ka), and Quelguenco and Chihuío at 39.9°S (K-Ar age of ca. 0.7 Ma; 40Ar/ 39Ar age of ca. 720 ka) are stratocones with well-preserved necks or radial dike swarms located at the Andean drainage divide (Lara et al., 2001). Huanquihué volcano has a Holocene pyroclastic cone over the northern flank, which shows the persistence of the magmatic activity. Other central vents like Mencheca at 40.5°S (K-Ar age of ca. 0.53 Ma), Cordón de Alvarez at 40.6°S, Fiuchá at 40.8°S, and Sarnoso at 40.8°S (K-Ar age of ca. 0.9 Ma) can be recognized at the base of Puyehue and Casablanca active volcanoes. Near 41°S, La Picada stratocone is located between the Quaternary Osorno and Puntiagudo volcanoes. The best-preserved stratovolcanoes in this group can have middle Pleistocene lavas that overlap the basal parts of the active stratovolcanoes.
10447-01
10449-01
10446-01
10359-01
XG-124
XG-102
XG-185
XG-175
Chihuío volcano
Pirihueico volcano
Sierra de Quinchilca
Nevados de Caburgua
Geological unit
39°55’S/71°36’W
39°58'S/71°37'W
39°40'S/72°00'W
39°10'S/71°32'W
Lat/Long (°)
0.0023014 ± 0.0000144
0.0012416 ± 0.0000489
0.0012700 ± 0.0000500
0.0012755 ± 0.0000502
J
GM
GM
GM
GM
3–30
3-20
3–16
3–22
Material Pw/°C
100.00
720 ± 30
623 ± 16
280 ± 90
93.40 76.10
340 ± 80
984 ± 11
Age (ka) ±2
100.00
85.16
Ar %
39
1.50
1.66
MSWD
0.29
Sums (N-2)
292.6 ± 1.2
Ar/36Ari ±2
40
5 of 5
5 of 6
1.77
0.93
296.2 ± 2.4
306.7 ± 1.9
3 of 5 (Weighted mean plateau)
5 of 5
4 of 6
N
Age (ka) ±2
716 ± 38
601 ± 20
1004 ± 16
Isochron Analysis
Note: GM—groundmass material free of phenocrysts, J: Irradiation parameter; Pw/ºC: Power/Temperature ratio. Bold for preferred ages. Analysis at SERNAGEOMIN (Servicio Nacional de Geología y Minería, Chile), calculated relative to 28.03 Ma Fish Canyon sanidine (Renne et al., 1994).
Experiment
Sample
Age Spectrum
TABLE 1. SUMMARY of 40Ar/39Ar INCREMENTAL HEATING EXPERIMENTS
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304
L.E. Lara and A. Folguera TABLE 2. PUBLISHED K-Ar AGES FOR PLIOCENE-QUATERNARY ROCKS (38°–42°S): SOUTHERN ANDES Sample
Site
UTM N UTM E Material
%K
Pliocene-Quaternary Western Volcanoes and Volcanic Sequences P253 Reloncaví s/n Pucón EL-790 Malleco 5764.5 254.15 WR 0.835 EL-339 Malleco 5756.2 284.075 WR 0.549 EL-646 Malleco 5779.75 260.65 WR 1.009 EL-841 Malleco 5752.8 254.1 WR 0.729 EL-666 Malleco 5789.13 245.8 WR 0.17 EL-920 Malleco 5691.8 239.98 WR 0.867 EL-334 Malleco 5755.25 289.5 WR 0.662 EL-644 Malleco 5781.65 260.9 WR 1.152 EL-728 Malleco 5789.85 268 WR 1.0101 EL-726 Malleco 5789.85 268 WR 0.053 EL-582 Malleco 5770 241.7 WR 1.854 EL-798 Malleco 5766.7 252.4 WR 1.924 XG-223A Malleco 5656.2 243.2 WR 0.368 XG-167 Maichin 5641.1 280.2 WR 0.639 XG-136 Trautrén 5663.9 288.7 WR 2.326 XG-138 Trautrén 5660.4 290 WR 0.618 XG-125 Caburgua 5662.4 280.2 WR 0.611 XG-124 Caburgua 5660.6 278.9 WR 1.331 XG-005 Cordillera Nevada 5527.4 772.5 WR 0.92 XG-070 Cordillera Nevada 5523.5 731.5 WR 1.045 XG-026 Cordillera Nevada 5522.2 715.7 WR 0.555 XG-029 Cordillera Nevada 5514.5 710.4 WR 0.907 XG-071 Cordillera Nevada 5521.5 730.3 WR 1.2 XG-102B Quinchilca 5603.2 756.4 WR 0.452 XG-104A Quinchilca 5602.3 756.1 WR 0.648 XC-205 Quinchilca 5670.2 276.7 WR 1.036 HM-73 Mencheca WR 1.08 XA-300 Sarnoso 5482.2 718.8 WR 1.02 XG-185 Pirihueico 5577.4 277.8 WR 0.674 XG-175 Chihuío 5572.5 278.5 WR 1.684 170185-2 La Picada 5560.7 712.1 WR 0.618 XC-288 Chapuco 5467.2 719.2 WR 1.01 XC-289 Chapuco 5462.6 719.2 WR 0.698 HO-46 Chapuco 5463 710.8 WR 0.709 HO-46 Chapuco 5463 710.8 WR 0.709 60385 Hueñu-Hueñu 5424 706 WR 0.48 090185-3 Reloncaví 5408.2 725.8 WR 0.507 P273b Cuernos del Diablo 5419.8 745.5 WR 0.92
The eastern belt of late Pliocene to early Pleistocene volcanoes occurs on the Copahue–Pino Hachado uplift on the west side of the Loncopué trough. The Bonete volcano at 37.4°S (K-Ar age of ca. 3.6 Ma) and the Trocolán volcano at 37.6°S are near the drainage divide. The Caldera del Agrio (K-Ar ages of ca. 4.3–0.8 Ma), which encloses the active Copahue volcano, is located at 37.9°S, where the Copahue–Pino Hachado block merges with the Main Cordillera. The Caldera del Agrio is part of a NE-trending, 80-km-long volcanic chain that includes the Mandolegüe volcano at 37.8°S, a monogenetic lava field (K-Ar ages of ca. 1.6–0.8 Ma), the Bayo and Trolope dome complexes (K-Ar of ca. 0.6 Ma), and the eroded Trolón (K-Ar of ca. 0.6 Ma) stratovolcano at 37.6°S (Pesce, 1989; Linares et al., 1999; Folguera et al., 2004). Stratovolcanoes farther south include the Las Monjas at 38.1°S, the Rahue (K-Ar ages of ca. 1.4–1.0 Ma) at 38.2°S, the Butahuao (ca. 1.5 Ma) at 38.3°S, the Pino Solo (K-Ar age of ca. 2.1 Ma) at 38.5°S, and the Tralilhue (K-Ar age of ca. 2.3 Ma) at 38.5°S (Muñoz and Stern, 1988). Partially
Vol.40Ar rad (nL/g)
0.074 0.039 0.065 0.049 0.96 0.133 0.091 0.072 0.062 1.034 0.063 0.06 0.024 0.023 0.07 0.013 0.057 0.046 0.049 0.049 0.018 0.021 0.019 0.014 0.041 0.011 0.034 0.038 0.049 0.01 0.017 0.015 0.027 0.023 0.027 0.005 0.026
% Ar atm.
Age + 2s
Reference
90 93 90 86 75 79 96 89 85 94 91 88 97 95 94 95 89 83 93 89 97 95 93 97 96 97 91 82 95 88 91 85 93 88 90 79 95 89
3.1 + 2.1 3.7 + 1.5 2.3 + 0.8 1.8 + 1.1 1.7 + 0.6 1.7 + 0.4 4.4 + 0.5 3.9 + 0.7 3.5 + 2.1 1.6 + 0.7 1.5 + 0.5 1.3 + 0.6 0.9 + 0.3 0.8 + 0.3 1.7 + 1.1 0.9 + 0.7 0.8 + 0.4 0.5 + 0.3 2.4 + 0.5 0.8 + 0.5 1.4 + 0.6 1.2 + 0.3 0.7 + 0.5 0.9 + 0.3 0.4 + 0.3 0.8 + 0.6 1.4 + 0.6 0.5 + 0.5 0.53 + 0.44 0.9 + 0.1 1.5 + 0.7 0.7 + 0.2 0.52 + 0.2 0.4 + 0.1 0.6 + 0.5 1 + 0.3 0.9 + 0.3 1.4 + 0.2 0.27 + 0.14 0.7 + 0.4
SNGM-BRGM (1995) SNGM-BRGM (1995) Suárez and Emparán (1997) Suárez and Emparán (1997) Suárez and Emparán (1997) Suárez and Emparán (1997) Suárez and Emparán (1997) Suárez and Emparán (1997) Suárez and Emparán (1997) Suárez and Emparán (1997) Suárez and Emparán (1997) Suárez and Emparán (1997) Suárez and Emparán (1997) Suárez and Emparán (1997) Lara et al. (2001) Lara et al. (2001) Lara et al. (2001) Lara et al. (2001) Lara et al. (2001) Lara et al. (2001) Campos et al. (1998) Campos et al. (1998) Lara et al. (2001) Lara et al. (2001) Lara et al. (2001) Lara et al. (2001) Lara et al. (2001) Lara et al. (2001) Moreno (1977) Lara et al. (2001) Lara et al. (2001) Lara et al. (2001) Moreno et al. (1985) Lara et al. (2001) Lara et al. (2001) SNGM-BRGM (1995) Lara et al. (2001) Moreno et al. (1985) Moreno et al. (1985) SNGM-BRGM (1995) (continued)
overlapping them are the basaltic lava flows from the Pino Hachado caldera (K-Ar ages from ca. 2.0 to 1.4 Ma) at 38.6°S, which has young pyroclastic cones in its interior. At the southern end of the belt is the Palao Mahuida caldera at 38.7°S and the Queli Mahuida volcano (ca. 1.0 Ma) at 38.8°S. East of the Agrio fold-and-thrust belt, the Tromen volcano is a prominent stratocone that has basal flows that have been 40Ar/ 39Ar dated between ca. 2.08 and 0.18 Ma (Galland, 2004; Kay et al., this volume, chapter 2). Farther east, Auca Mahuida volcano is an extended shield where Rossello et al. (2002) and Kay et al. (this volume, chapter 2) have reported 40Ar/ 39Ar ages ranging from ca. 1.78 to ca. 0.88 Ma. Middle Pleistocene Middle Pleistocene centers are nearly indistinguishable from late Pliocene to early Pleistocene centers with respect to their morphology and extent of erosion. Many of these volca-
Pliocene to Quaternary narrowing of the Southern Andean volcanic arc
305
TABLE 2. PUBLISHED K-Ar AGES FOR PLIOCENE-QUATERNARY ROCKS (38°–42°S): SOUTHERN ANDES (continued) Sample
Site
UTM N UTM E Material
%K
Pliocene-Quaternary Eastern Volcanoes and Volcanic Sequences TC-31 Copahue WR 1.866 TC-36 Copahue WR 6.78 CO-36 Copahue WR 1.79 CO-37 Copahue WR 2.88 CO-39 Copahue WR 2.34 CO-40 Copahue WR 2.35 CO-7 Copahue WR 2.51 CO-29 Copahue WR 2.65 CO-49 Copahue WR 1.67 CO-52 Copahue WR 2.61 CO-19 Copahue WR 3.53 CO-20 Copahue WR 3.5 Ba-1 Copahue WR 3.22 CC-99 Copahue WR 0.91 CO-2 Hualcupén WR 1.03 CO-2 Hualcupén WR 1.03 CO-10 Hualcupén WR 1.41 CO-10 Hualcupén WR 1.41 CO-23 Hualcupén WR 1.19 CO-23 Hualcupén WR 1.19 CO-32 Hualcupén WR 1.27 CO-47 Hualcupén WR 1.8 CO-48 Hualcupén WR 1.65 CO-48 Hualcupén WR 1.65 AG-1 Hualcupén WR 1.18 AG-1 Hualcupén WR 1.18 CO-22 Mellizas WR 1.7 CO-33 Mellizas WR 1.82 CO-34 Mellizas WR 2.58 CO-1 Mellizas WR 3.51 CO-3 Mellizas WR 3.46 CO-3 Mellizas bt 6.76 TC-78 Rahue 5733.25 322.25 WR 1.461 TC-84 Rahue 5768.25 329.75 WR 2.102 TC-80 Rahue 5763 318 WR 2.501 TC-75 Rahue 5764.75 323 WR 1.864 CZ-37 Pino Solo 5734.5 337.5 WR 2.296 CZ-40 Pino Solo 5733.25 335.5 WR 2.96 CZ-23 Pino Hachado 5721.9 318 WR 3.189 CZ-30 Pino Hachado 5726.25 333.75 WR 2.495 TC-61 Pino Hachado WR 1.348 TC-62 Pino Hachado 5721.9 325.67 WR 2.39 CZ-32 Pino Hachado 5720.5 335.25 WR 2.65 TC-55 Queli Mahuida WR 1.364 TC-57 Queli Mahuida WR 1.675 EL-460 Trubul 5769 302.43 WR 0.797 TC-87 Huisa 5768.5 313.75 WR 1.043 TC-95 Cayulafquen 5742.5 318 WR 1.893 TC-119 Tralihue 5738 329.75 WR 2.791 CZ-51 Tralihue 5736.5 333 WR 2.196 IGE-9 Bateamahuida 5703 310.25 WR 0.703 IGE-8 Bateamahuida 5701.73 309.15 WR 0.754 EL-180 Ranquil 5762.55 303.025 WR 0.944 EL-459 Ranquil 5768.75 301.875 plag 0.406 IGE-327 Ranquil 5709.73 318.5 WR 1.162 EL-421 Ranquil 5765.25 300.45 plag 0.356 EL-400 Ranquil 5770.25 299.2 plag 0.256 EL-397 Ranquil 5768.75 299.25 plag 0.375 1 WR 0.607 2 WR 0.652 3 WR 0.876 XM-23 Tronador (GDU) WR 1.85 Note: WR—whole rock, bt—biotite, plag—plagioclase.
Vol.40Ar rad (nL/g)
% Ar atm.
Age + 2s
Reference
0.06 0.06 0.036 0.038 0.058 0.035 0.071 0.05 0.43 0.037 0.055 0.067 0.034 0.01 0.076 0.079 0.093 0.105 0.105 0.091 0.125 0.133 0.119 0.136 0.102 0.084 0.078 0.082 0.12 0.127 0.188 0.241 0.059 0.24 0.402 0.101 0.188 0.254 0.193 0.142 0.075 0.151 0.169 0.054 0.057 0.037 0.042 0.128 0.264 0.196 0.142 0.144 0.177 0.066 0.204 0.061 0.032 0.049 0.032 0.025 0.058 0.095
82 94 96 84 28 76 83.4 79.3 88 93.3 77.5 68.8 94.5 95.6 80.3 76.6 58.2 59.7 69.9 55.8 73.1 92 71.3 22.7 52.8 78.3 62.2 42.4 77.4 77.8 54 90.6 85 89 64 92 62 82 83 79 76 75 76 88 93 96 92 85 75 58 72 74 78 96 78 90 86 87 93 91 76 94
0.8 + 0.1 1.1 + 0.5 1.16 + 0.36 0.76 + 0.28 1.23 + 0.36 0.91 + 0.28 1.63 + 0.1 1.09 + 0.1 1.48 + 0.14 0.82 + 0.16 0.9 + 0.14 1.1 + 0.18 0.62 + 0.12 0.66 + 0.14 4.25 + 0.1 4.42 + 0.14 4 + 0.1 4.29 + 0.1 5.08 + 0.14 4.4 + 0.1 5.67 + 0.14 4.26 + 0.1 4.15 + 0.1 4.75 + 0.1 4.98 + 0.14 4.1 + 0.1 2.64 + 0.08 2.6 + 0.1 2.68 + 0.14 2.08 + 0.16 2.63 + 0.2 2.05 + 0.1 1 + 0.2 2.9 + 0.5 4.1 + 0.3 1.4 + 0.4 2.1 + 0.2 2.2 + 0.3 1.6 + 0.2 1.5 + 0.1 1.4 + 0.2 1.6 + 0.2 1.6 + 0.1 1 + 0.4 0.9 + 0.3 1.2 + 0.7 1 + 0.5 1.7 + 0.3 2.4 + 0.3 2.3 + 0.2 5.2 + 0.6 4.9 + 0.4 4.8 + 0.5 4.1 + 2.5 4.5 + 0.5 4.4 + 2 3.2 + 2 3.3 + 2 1.4 + 0.4 1 + 0.3 1.7 + 0.2 1.3 + 0.3
Muñoz and Stern (1988) Muñoz and Stern 1(988) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Linares et al. (1999) Muñoz and Stern (1988) Muñoz and Stern (1988) Muñoz and Stern (1988) Muñoz and Stern (1988) Muñoz and Stern (1988) Muñoz and Stern (1988) Muñoz and Stern (1988) Muñoz and Stern (1988) Muñoz and Stern (1988) Muñoz and Stern (1988) Muñoz and Stern (1988) Muñoz and Stern (1988) Muñoz and Stern (1988) Suárez and Emparán (1997) Muñoz and Stern (1988) Muñoz and Stern (1988) Muñoz and Stern (1988) Muñoz and Stern (1988) Suárez and Emparán (1997) Suárez and Emparán (1997) Suárez and Emparán (1997) Suárez and Emparán (1997) Suárez and Emparán (1997) Suárez and Emparán (1997) Suárez and Emparán (1997) Suárez and Emparán (1997) Folguera et al. (2004) Folguera et al. (2004) Folguera et al. (2004) Mella et al. (2005)
306
L.E. Lara and A. Folguera TABLE 3. 40Ar/39Ar AGES FOR PLIOCENE-QUATERNARY VOLCANOES (38°–42°S): SOUTHERN ANDES Sample
Geological unit
Tromen volcano ABL-05 Tilhué Formation ABL-05 Tilhué Formation ABL-04 Dikes ABL-06 Sill ABL-07 Dikes ABL-03-09 Dikes TR-01 Pichichaico Formation TR-02 Pichichaico Formation CB-03-12 Tilhué Formation CB-03-01 Tilhué Formation CB-03-04 Tilhué Formation CB-03-06 Tilhué Formation CB-03-13 Tilhué Formation CB-03-14 Tilhué Formation LB-01 Basalto III Formation LB-02 Basalto III Formation TR-03 Basalto IV Formation YS-03-02 Basalto III Formation LB-03 Basalto V Formation TIL-01 Tilhué Formation TIL-04 Tilhué Formation TIL-06 Tilhué Formation TIL-07 Tilhué Formation Tronador volcano XB-29 XB-32 XM-7 XM-22
Tronador II Tronador II Tronador III Tronador III
Auca Mahuida ASR1 AM1 AM3 AM4 AM5 AM6 AM8 AM9 AM10 AM11
Material
Age (ka) ±2s
39 Ar (%)
Reference
biotite biotite WR WR WR WR WR WR WR biotite biotite WR biotite biotite WR WR WR WR WR biotite WR biotite biotite
2.01 ± 0.52 2.08 ± 0.26 1.98 ± 0.16 1.95 ± 0.26 2.03 ± 0.10 1.61 ± 0.32 1.81 ± 0.12 1.81 ± 0.12 1.72 ± 0.04 0.9 ± 0.08 1.1 ± 0.14 1.1 ± 0.12 1.21 ± 0.16 0.84 ± 0.18 1.25 ± 0.14 1.3 ± 0.24 1.8 ± 0.08 1.75 ± 0.12 0.99 ± 0.30 0.83 ± 0.41 0.88 ± 0.30 0.82 ± 0.08 1.44 ± 0.25
95.0 98.0 96.0 23.0 99.9 96.0 100.0 100.0 99.5 79.0 99.0 48.0 96.0 97.0 99.0 99.0 97.0 99.0 90.0
Galland (2004), unpublished Galland (2004), unpublished Galland (2004), unpublished Galland (2004), unpublished Galland (2004), unpublished Galland (2004), unpublished Galland (2004), unpublished Galland (2004), unpublished Galland (2004), unpublished Galland (2004), unpublished Galland (2004), unpublished Galland (2004), unpublished Galland (2004), unpublished Galland (2004), unpublished Galland (2004), unpublished Galland (2004), unpublished Galland (2004), unpublished Galland (2004), unpublished Galland (2004), unpublished Galland (2004), unpublished Galland (2004), unpublished Galland (2004), unpublished Galland (2004), unpublished
WR WR WR WR
0.53 ± 0.13 0.47 ± 0.04 0.36 ± 0.05 0.34 ± 0.04
Mella et al. (2005) Mella et al. (2005) Mella et al. (2005) Mella et al. (2005)
WR WR WR WR WR WR WR WR WR WR
1.4 ± 0.30 1.18 ± 0.04 1.61 ± 0.07 1.10 ± 0.20 1.19 ± 0.05 0.88 ± 0.03 1.70 ± 0.20 0.90 ± 0.07 1.16 ± 0.04 1.26 ± 0.07
Rosello et al. (2002) Rosello et al. (2002) Rosello et al. (2002) Rosello et al. (2002) Rosello et al. (2002) Rosello et al. (2002) Rosello et al. (2002) Rosello et al. (2002) Rosello et al. (2002) Rosello et al. (2002)
noes are partially collapsed structures that have upper units older than ca. 200 ka, and consequently are older than the basal units of the Pleistocene to Holocene centers. Their basal units can be early Pleistocene in age. Among these centers is the undated Quinquilil or Colmillo del Diablo volcano at 39.5°S, which consists of a prominent neck surrounded by basaltic lavas. Others are the Hawaiian-type centers at Sierra de Quinchilca at 40°S (K-Ar age of ca. 1.4–0.8 Ma; a 40Ar-39Ar age of ca. 340–280 ka) and Cordillera Nevada at 40.5°S. The Mirador volcano at 40.5°S is an eroded stratocone. The Pantoja volcano at 40.1°S is a prominent neck surrounded by basaltic lava flows (Fig. 2). The Cordillera Nevada caldera is a long-lived center with a base that has an age of ca. 500 ka (Lara et al., 2002) and an evolution that overlaps that of the Cordón Caulle–Puyehue volcanic complex (Moreno, 1977; Campos et al., 1998). Late Pleistocene–Holocene: The Present “Narrow” Volcanic Arc Volcanoes in the active arc form this group. Most of them are in the Main Cordillera on the western side of the Andes where they define a clear volcanic front that is ~250 km east of
96.0 94.0 84.0
the trench. Several volcanic complexes and many monogenetic Holocene cones occur along the intra-arc Liquiñe-Ofqui fault system (Fig. 1). Other stratocones to the east occur along oblique pre-Andean structures (Fig. 1). Together, these centers define an arc region with a maximum width of ~60 km. Farther east, retroarc Holocene pyroclastic cones like Laguna Blanca and Loncoloán occur in the southern part of the Loncopué trough. Except for the Tromen volcano near 37.3°S, Quaternary stratovolcanoes are not found in the extra-Andean region. At Tromen, Llambías et al. (1982) described possible Holocene flows and Kay et al. (this volume, chapter 2) reported a 40Ar/ 39Ar age of 0.175 ± 0.4 Ma for a young flow with a very weak arc geochemical signatures. The few available ages for the glacially eroded basal units of the active stratovolcanoes south of 38°S are younger than ca. 200 ka (Drake et al., 1976). Moreno et al. (1986) published K-Ar ages of ca. 171–146 ka for the basal units of the Callaqui volcano (37.9°S). Similar ages were obtained for the Osorno (K-Ar age of ca. 150 ka; Moreno et al., 1985) and Calbuco volcanos (K-Ar age of ca. 110 ka; Moreno et al., 1985). Others, like the Lanín volcano (Lara, 2004; Lara et al., 2004b, and references therein) and the Cordón Caulle–Puyehue volcanic com-
Pliocene to Quaternary narrowing of the Southern Andean volcanic arc
307
Figure 2. Photographs of representative morphologies of some Pliocene-Quaternary volcanoes from the western and eastern belts. Ages are indicated where available. (A) Pirihueico volcano (39.9°S) and radial dikes that intrude upper lavas (ca. 601 ka) at the drainage divide of the Andes; (B) neck of Pantoja volcano (40.1°S) surrounded by basaltic lavas; (C) basal lavas (ca. 500 ka) from La Picada volcano (41.1°S) at the present volcanic front; (D) Mandolegüe volcano (37.8°S), north of Caldera del Agrio; (E) Pino Solo volcano (38.5°S) on the Copahue–Pino Hachado block; and (F) Guañaco neck (37.3°S) on the eastern Andean foothills.
plex (Lara et al., 2002), are younger than ca. 100 ka. Some stratovolcanoes have mixed ages, whereas others are strictly late Pleistocene or entirely Holocene.
tical approach. Four new 40Ar/ 39Ar ages in Table 1 from eroded centers at the present arc front or the Andean water divide are also discussed.
K-Ar AND 40Ar/ 39Ar GEOCHRONOLOGY
Analyses of Published K-Ar Ages
A key problem in discerning the architecture and evolution of the late Cenozoic volcanic arc in the Southern Andes is the lack of sufficient reliable geochronologic data. The basaltic low-K magmas that predominate in the volcanic centers in the arc segment between 38° and 41°S lack suitable phenocrysts for K-Ar or 40Ar/ 39Ar dating. As a first approach, the existing whole-rock K-Ar ages are evaluated here using a simple statis-
As has been widely discussed, K-Ar whole-rock ages can be less precise than 40Ar/ 39Ar ages (e.g., Singer et al., 1997). Nevertheless, they avoid some of the obstacles encountered in irradiating fine-grain or glassy samples prepared for 40Ar/ 39Ar (Lanphere, 2000). The main part of the K-Ar data set here (103 analyses, Table 2) comes from Lara et al., (2001), Muñoz and Stern (1988), Suárez and Emparán (1997), Linares et al. (1999),
308
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and recent geological maps of the Servicio Nacional de Geología y Minería (Chile) and references therein. A statistical analysis of the K-Ar ages for the PlioceneQuaternary western arc volcanoes shows a main peak of relative probability (Deino and Potts, 1990) at ca. 1.0 Ma and a minor one at ca. 2.6 Ma (Fig. 3). The ca. 2.6 Ma peak comes entirely from the Copahue volcanic complex that forms part of both the western and the eastern volcanic domains. Younger ages from the inner retroarc area show two relative peaks at ca. 0.83 and ca. 1.8 Ma. Although the main peaks could indicate a sampling bias for a small population (n = 103 samples), this result supports a constant Cenozoic magmatism in the arc front and nearly coeval volcanism in the inner retroarc. New 40Ar-39Ar Ages New 40Ar/ 39Ar ages were obtained from four samples from morphologically recognizable centers in the western PlioceneQuaternary volcanic belt. All were from upper units, and thus the underlying basal flows must be older. Age spectra are shown in Figure 3, and analytical methods are discussed in Appendix 1. The first sample, a silicic andesite (59.9% SiO2; 8% plagioclase [plag] + 2% clinopyroxene [cpx]) from the Nevados de Caburgua (39°S) complex, for which Lara et al. (2001) obtained a K-Ar age of 0.8 ± 0.5 Ma, gave a plateau age of 984 ± 11 ka with 85.2% of the released gas. A second sample, a basaltic andesite (54.6% SiO2; 25% plag + 5% cpx) from Chihuío volcano (39.5°S), for which Lara et al. (2001) obtained a K-Ar age of 0.7 ± 0.2 Ma, gave a plateau age of 720 ± 30 ka with 100% of released gas. A basalt (49.1% SiO2; 25% plag + 12% olivine [ol] + 3% cpx) from the Pirihueico volcano, for which Lara et al. (2001) obtained a K-Ar age of 1.5 ± 0.7 Ma, gave an 40Ar/ 39Ar isochron age of 601 ± 20 ka for 76.1% degassing. A slight Ar excess was detected in the step-heating experiment. Finally, a basaltic andesite (52% SiO2; 10% plag + 3% cpx + 2% ol) from the Sierra de Quinchilca (39.5°S) caldera, for which Lara et al. (2001) obtained a K-Ar age of 0.8 ± 0.6 Ma, gave a plateau age of 340 ± 80 ka. This sample yielded a preferred weighted mean age of 280 ± 90 ka for the three first steps with 93.4% of gas releasing. An Ar excess was recognized in the last steps. GEOCHEMISTRY With some exceptions, samples from Pliocene-Quaternary volcanoes mimic the first-order geochemical features of the active Southern Andean stratovolcanoes. Only a slight departure in space and time is observed. The graphs in Figures 4A to 4D show temporal and spatial cross-arc comparisons of some key chemical characteristics. Samples from the eastern Pliocene centers show strong alkali enrichment at low silica contents (Fig. 4A). These samples define a high-alkali trend, as do the Pliocene-Quaternary samples from the Pino Hachado caldera (Muñoz and Stern, 1988, 1989).
In contrast, western Pliocene and Pliocene-Quaternary samples overlap the field of Quaternary arc-front samples (Villarrica, Osorno, and Puyehue; Gerlach et al., 1988; Hickey-Vargas et al., 1989; Tagiri et al., 1993). Interestingly, Pliocene-Quaternary samples from the Copahue–Pino Hachado block show a trend like the modern easternmost-arc center samples (Lanín; Lara et al., 2004b), which is 80 km to the south. Samples from PlioceneQuaternary centers (ca. 1.3–0.3 Ma) at the present drainage divide coincide with the overall Pliocene-Quaternary suite. Ba/La ratios, which are commonly related to the input of slab-derived fluids into the mantle source, decrease to the east in both Pliocene-Quaternary and Quaternary rocks (Fig. 4B). Decreasing Ba/La ratios associated with increasing light rare earth elements (LREEs) (La/Sm) can be interpreted as indicating lower degrees of partial melting in the mantle to the east (e.g., Hickey-Vargas et al., 1989). As with alkali contents, eastern Pliocene-Quaternary samples plot in the same field as the Quaternary arc samples at the drainage divide. The Laguna Blanca samples from the Loncopué trough have low Ba/La ratios, showing a backarc signature. La/Nb ratios show a similar trend to Ba/La ratios and are consistent with an eastward decrease in slabderived components in the mantle wedge (Fig. 4C). Light REE to high field strength element ratios (La/Nb) in the Pino Hachado caldera and Laguna Blanca basalts overlap the field of oceanicisland basalts (OIBs) and are only slightly more enriched than intraplate basalts of southern Patagonia (Stern et al., 1990). REE data show the same relative regional patterns (Fig. 4D). A sharp east-west LREE enrichment can be observed among both the Pliocene-Quaternary and Quaternary samples. The magmas of all of these centers have been argued to have experienced low-pressure fractional crystallization in a thin crust on the basis of their differentiation trends and a lack of evidence for significant interaction with the continental crust in their Sr and Nd isotopic ratios (Muñoz and Stern, 1989). The lowest degrees of melting is inferred for the Laguna Blanca backarc basalts based on their relatively high LREE contents and steep heavy (H) REE patterns (Fig. 4D). The trachydacitic magmas (ca. 1.6 Ma) from the Pino Hachado volcanic complex show an end member–like geochemical signature (Fig. 4A–D), which departs from that of the coeval eastern Pliocene-Quaternary trends. Their more alkaline affinities are shown by high alkali contents at a given silica content, low Ba/La and La/Nb ratios, and LREE enrichment. These features are consistent with a low degree of melting of a backarc-type asthenospheric source. TECTONIC EVOLUTION Whereas the first-order features of the volcanic arc can be related to plate interactions at a continental scale, the local stress regime of the upper crust seems to play a direct role in the spatial distribution of the volcanic centers along the arc. In the Main Cordillera, the Pliocene-Quaternary volcanic centers are located along the margin-parallel Liquiñe-Ofqui fault system
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Figure 3. K-Ar age distribution and relative probability graphs (Deino and Potts, 1990) for Pliocene-Quaternary centers from the western and eastern belts (upper panel). Black circles represent ages from the western Pliocene-Quaternary belt (mainly from Lara et al., 2001; Suárez and Emparán, 1997, and references therein). Ages from the North Patagonian Batholith (NPB) and Pliocene volcanic sequences are also included (mainly from Lara and Moreno, 2004; Rodríguez et al., 1999; Campos et al., 1998; Cembrano et al., 1996, and references therein). Gray circles represent ages from the eastern Pliocene-Quaternary belt (mainly from Linares et al., 1999; Muñoz and Stern, 1988; Rossello et al., 2002; and Galland, 2004). Diamonds show selected 40Ar/ 39Ar ages. Square is for U-Pb age from the North Patagonian Batholith. Lower panel shows degasification diagrams for selected samples from the western Pliocene-Quaternary belt. Plateau and isochron ages are at a 2σ level of uncertainty.
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Figure 4. (A) Alkali versus silica contents; (B) Ba/La versus La/Sm ratios; (C) La/Nb versus La and; (D) rare earth element (REE) abundances for Pliocene-Quaternary volcanoes. Data are from Lara et al. (2001), Suárez and Emparán (1997), and Muñoz and Stern (1988). Values for Quaternary volcanoes are mainly from Lara et al. (2004b); Mella et al. (2005); Tagiri et al. (1993); Hickey-Vargas et al. (1989); and Gerlach et al. (1988). Normalizing factors for REEs were taken from Gerlach et al. (1988) and references therein (La 0.315; Ce 0.813; Pr 0.096; Nd 0.597; Sm 0.192; Eu 0.072; Gd 0.2043; Tb 0.049; Dy 0.254; Ho 0.0567; Er 0.166; Yb 0.208; Lu 0.02539). OIB—oceanic-island basalt.
and NW-trending structures. To the east, the Pliocene-Quaternary volcanoes occur in pre-Andean blocks and along the main faults that bound these uplifted blocks (Fig. 1). From Miocene to Pliocene times, high-strain domains were established along the Liquiñe-Ofqui fault system, which controlled pluton ascent in a dextral transpressional regime (Cembrano et al., 2000, 2002). In the foreland to the east, mild late Miocene deformation in the Agrio fold-and-thrust belt produced out-of-sequence reactivations of extensional structures that had
been previously inverted in the Late Cretaceous. These structures are associated with synorogenic strata that accumulated mainly in the inner retroarc area (Ramos, 1998; Zapata et al., 1999). Brittle deformation along the Main Cordillera records an E-W contractional event that lasted until the late Pliocene (Lavenu and Cembrano, 1999a, 1999b). A phase of orogenic relaxation occurred during the early Pliocene in the eastern foothills of the Andes, with the emplacement of volcaniclastic sequences with main depocenters that were controlled by exten-
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Figure 4 (continued).
sional faults (Folguera et al., 2002, 2003). The Quaternary stress regime shows dextral transpression along the volcanic arc, as indicated by microstructure tectonic data (Lavenu and Cembrano, 1999a, 1999b) and scarce focal mechanisms of southern Andean crustal earthquakes (e.g., Barrientos and Acevedo, 1992; Chinn and Isacks, 1983). The time of change is not clear, but Lavenu and Cembrano (1999a, 1999b) suggested a minimum age of ca. 1.6 Ma based on 40Ar/ 39Ar ages of mylonites that locally predate Quaternary brittle deformation. Postorogenic collapse would explain transient dilatation of the overall structure, especially in the inner retroarc region.
times (Kay et al., 2005). Subduction of a younger Nazca plate beneath a thin continental crust and the presence of a marginscale intra-arc fault system from the triple junction (46°S) to the northern boundary of the segment (38°S) account for a stable magmatic front since at least the Miocene. South of 38°S, no arc migration or steepening of the slab can be inferred from the available data. From the late Pliocene, changes in magmatism are expressed in volume of magma production and geochemical variations across the arc.
DISCUSSION
South of 38°S, Muñoz and Stern (1988) proposed that the arc front migrated to the west in the Pleistocene based on the presence of late Pliocene to early Pleistocene ages in Copahue–Pino Hachado block volcanic rocks, and the absence of such ages in volcanic rocks along the Main Cordillera. Since
South of 38°S, subduction geometry, magmatism, and tectonic regimes do not record the dramatic changes inferred for the northernmost southern Andes (33°–38°S) during Cenozoic
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that time, new K-Ar ages and the new 40Ar/ 39Ar ages in Table 1 for Pliocene-Quaternary volcanoes along the western foothills of the Andes show that this frontal arc shift did not occur and that Pliocene-Quaternary volcanoes of the Main Cordillera are underlain by volcanic rocks with K-Ar ages from ca. 2.4 Ma to ca. 5.2 (e.g., Lara and Moreno, 2004). Intrusives in the North Patagonian Batholith have U-Pb crystallization ages as young as 5.3 ± 0.8 Ma (e.g., Campos et al., 1998). South of 41°S, where exhumation rates are higher than to the north (Gräfe et al., 2002), granodiorites in the Main Cordillera near the Liquiñe-Ofqui fault system have late Pliocene 40Ar/ 39Ar mineral ages as young as 3.4 ± 0.5 and 3.6 ± 0.3 Ma (Cembrano et al., 2002). Mylonites with ages of 3.59 ± 0.01 and 3.78 ± 0.01 Ma along the fault are next to undeformed plutons (Cembrano et al., 2000, 2002). Given that ductile deformation was nearly coeval with pluton emplacement, the locus of the Pliocene magmatic front was the same in the Miocene. As such, the Quaternary volcanic front overlaps the Pliocene front, and the age pattern records a Pliocene to Quaternary narrowing, not a migration. Moreover, an arc migration hypothesis caused by slab steepening has problems in this region due to the inertial properties of the slab, which preclude a fast response to a decrease in the subduction angle in a million-year time scale (Turcotte and Schubert, 2002). An alternative suggested by Stern (1989) is accretion along the inner trench wall and westward migration of the trench. Although high rates of climatically driven erosion would cause rapid rates of sedimentary infilling of the trench, the active intracanyon currents would disperse these sediments to the north (Thornburg et al., 1990). In addition, negative rollback, as has been proposed for the Cenozoic evolution north of 38°S (Folguera et al., this volume, chapter 12; Kay et al., 2005), would cause trench retreat and a subsequent arc migration. South of 38°S, there is no clear evidence for trench retreat. These arguments along with the age data argue against westward migration of the volcanic front. As proposed by Molnar et al. (1979) and later refined by Shimozuru and Kubo (1983), a direct relation could exist between convergence velocities and widths of volcanic arcs. Invigorated dynamics of the subarc mantle is a plausible response to high subduction velocities, which could result in extended volcanic arcs with arc-front geochemical signatures far off the trench. Molnar et al. (1979) proposed an empirical relation between subduction velocity, the age of the subducted slab, and the length of the seismic Wadati-Benioff zone above the slab: L (km) ≈ V (mm/yr) × T (Ma). For a static volcanic front, L (length) can be expressed as a function of the forearc distance and the arc width. For a given subduction angle of ~35° (Cahill and Isacks, 1992; Bohm et al., 2002), we can rewrite this equation quantifying the effect of a decreasing velocity. At ~40°S and 2–3 Ma, the subduction velocity was ~9 cm/yr (Engebretson et al., 1986; De Mets et al., 1994) decreasing to 7.9 cm/yr (De Mets et al., 1994; Tamaki, 2000) or even to 5–6 cm/yr (Angermann et al., 1999) at present, which allows a maximum reduction in arc width of ~50%. For com-
parison, a slight steepening of the slab from 35° to 40° could cause an ~80% arc narrowing. Because the eastern belt is not parallel to the western Pliocene-Quaternary volcanic belt, a decelerating Nazca plate could play a role but cannot be the sole driving force for the arc narrowing. The presence of margin-scale intra-arc faults could help to modify the position of the magmatic front or the entire volcanic domain. They would act as suction pumps localizing fluids, which would result in a modified arc architecture. Arc narrowing, whatever the causal factors, was the process south of 38°S in the Pliocene to Quaternary transition. Time of Arc Narrowing The limited amount of Pliocene to Quaternary geochronologic data in the Southern Andes precludes a precise temporal definition of the time of arc narrowing. A statistical analysis of published K-Ar data and new 40Ar/ 39Ar ages allows a first approximation. Several points are clear. One is that the youngest eruptive units from the partially eroded centers at the static arc front are as old as ca. 0.5 Ma, and basal units must be older. Eroded centers from the arc are as old as ca. 1 Ma and overlap the age range of the basement sequences. Second, the uppermost lavas from the eroded centers at the modern drainage divide have arc signatures similar to the eastern Quaternary volcanoes (Lanín), which are as old as ca. 720 ka. Thus, the arc narrowing should have occurred before these lavas erupted; the opposite case would have caused them to have arc-front signatures. Further, eroded centers from the Copahue–Pino Hachado block with ages from ca. 4.1–0.8 Ma generally have arc chemical signatures like eastern Quaternary volcanoes, such as Lanín in the southern Andes. Excluding Copahue volcano, which has arc-front signatures as expected for its static pivotal position, Pino Hachado caldera (ca. 1.6–1.4 Ma) and Queli Mahuida volcano (ca. 1.0–0.9 Ma), the youngest members of the eastern suite (ca. 1.6–0.8 Ma), have near backarc signatures (Fig. 4). Interestingly, a trachyandesite from Pino Hachado caldera (ca. 1.6 Ma) has features that suggest the lowest degrees of mantle melting and the weakest slab input, similar to the neighboring Holocene Laguna Blanca basalts. Therefore, progressively from ca. 1.6 Ma, these eastern centers would have been in a backarc setting and would have recorded the narrowing of the arc. Coincidently, the beginning of the Quaternary dextral transpressional regime along the Southern Andean Main Cordillera was established at ca. 1.6 Ma (Lavenu and Cembrano, 1999a). CONCLUSIONS An analysis of published K-Ar and new 40Ar/ 39Ar ages, geochemical, and field tectonic data permit the following conclusions: (1) Nearly constant Late Cenozoic magmatism occurred along the Andean range in a magmatic arc front on the western side of the Main Cordillera. (2) South of 38°S, the PlioceneQuaternary volcanic arc had the same front as the modern one,
Pliocene to Quaternary narrowing of the Southern Andean volcanic arc ~250 km east of the trench. (3) Across-arc chemical variations of Pliocene-Quaternary magmas are similar to those in the Quaternary arc, but are displaced in space and time. (4) A westward arc narrowing, not a migration, occurred in the early Pleistocene. (5) A possible cause of arc narrowing was a decrease in subduction velocity 2–3 m.y. ago. (6) A plausible age for the beginning of arc narrowing is ca. 1.6 Ma, which is coincident with renewed dextral transpression along the arc. (7) The most active Quaternary volcanoes are spatially related to the Liquiñe-Ofqui fault system, suggesting that magma ascent will continue to concentrate in this region as has been the case in the Holocene. ACKNOWLEDGMENTS The 40Ar-39Ar analyses were done at the Servicio Nacional de Geología y Minería (SERNAGEOMIN, Chile) by C. Pérez de Arce and S. Mathews. Field work was partially supported by SERNAGEOMIN and Fondecyt grant No. 1960885. We acknowledge J. Muñoz and S.M. Kay for their comments. This paper is a contribution to the Volcanic Hazards Program of the SERNAGEOMIN and to the UNESCOIUGG-IGCP 455 project. APPENDIX 1: ANALYTICAL METHODS FOR 40Ar/ 39Ar DATING Samples for dating were crushed to a size between 250 and 80 μm and then handpicked to extract major phenocrysts and weathered surfaces. Single aliquots were analyzed by incremental heating with a CO2 laser at the SERNAGEOMIN (Servicio Nacional de Geología y Minería, Chile). Whole rocks were first placed in a disk of high-purity aluminum with a monitor grain of Fish Canyon sanidine (28.03 ± 0.1 Ma; Renne et al., 1994). Samples were irradiated in the La Reina nuclear reactor (Chile) for a period of ~48 h. Once the samples returned from the reactor, individual total fusion analyses were performed for all the monitors from the disk, and J factors were calculated for each grain, which represents an individual position in the disk. The distribution of J in two dimensions (2-D) across the disk was modeled by a 2-D quadratic fit to the data, resulting in a “J surface” for the disk (Arancibia et al., 2005). Individual J factors for each sample were calculated depending upon the coordinates of the sample (Table 1). Samples were analyzed by successive heating with increments of temperature by increases in the power of the CO2 laser with a maximum power of 30 W. Following each three heating steps, a line blank was analyzed. Then, the noble gases were separated from the other evolved gases by means of cold trap at –133 °C. Once purified, the noble gases were introduced into a high-resolution MAP 215–50 mass spectrometer in electron multiplier mode. The baseline was analyzed at the beginning and end of the analysis, for each step, and subtracted from the peak heights. Spectrometer bias was corrected using periodic analyses of air samples, from which a correction factor was calculated.
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Pliocene to Quaternary narrowing of the Southern Andean volcanic arc Ramos, V.A., 1998, Estructura del sector occidental de la faja plegada y corrida del Agrio, cuenca Neuquina, Argentina, in Proceedings, 10th Congreso Latinoamericano de Geología: Buenos Aires, Asociación Geológica Argentina, v. 2, p. 105–110. Renne, P.R., Deino, A.L., Walter, R.C., Turrin, B.D., Swisher, C.C., III, Becker, T.A., Curtis, G.H., Sharp, W.D., and Jaouni, A.R., 1994, Intercalibration of astronomic and radioisotopic time: Geology, v. 22, p. 783–786, doi: 10.1130/0091-7613(1994)022<0783:IOAART>2.3.CO;2. Rodríguez, C., Pérez, Y., Moreno, H., Clayton, J., Antinao, J., Duhart, P., and Martin, M., 1999, Area de Panguipulli-Riñihue, región de los Lagos: Santiago, Servicio Nacional de Geología y Minería, Mapas Geológicos no. 10, scale 1:100,000, 1 sheet. Rossello, E.A., Cobbold, P.R., Diraison, M., and Arnaud, N., 2002, Auca Mahuida (Neuquén Basin, Argentina): A Quaternary shield volcano on a hydrocarbon-producing substrate, in Proceedings, 5th International Symposium on Andean Geodynamics: Toulouse, IRD Editions, p. 549–552. SERNAGEOMIN-BGRM, 1995, Carta metalogénica X región sur: Servicio Nacional de Geología y Minería (SERNAGEOMIN), Bureau de Recherches Géologiques et Minières (BGRM) Open-File Report IR-95-05, v. 4, 10 sheets. Shimozuru, D., and Kubo, N., 1983, Volcano spacing and subduction, in Shimozuru, D., et al., eds., Arc volcanism: Physics and tectonics: Tokyo, Terra Publishers, p. 141–151. Singer, B.S., Thompson, R.A., Dungan, M.A., Feeley, T.C., Nelson, S.T., Pickens, J.C., Brown, L.L., Wulff, A.W., Davidson, J.P., and Metzger, J., 1997, Volcanism and erosion during the past 930 ky at the Tatara– San Pedro complex, Chilean Andes: Geological Society of America Bulletin, v. 109, no. 2, p. 127–142, doi: 10.1130/0016-7606(1997)109 <0127:VAEDTP>2.3.CO;2. Stern, C.R., 1989, Pliocene to Present migration of the volcanic front, Andean Southern volcanic zone: Revista Geológica de Chile, v. 16, no. 2, p. 145–162. 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
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Geological Society of America Special Papers Geochemistry and isotopic characteristics of the Caviahue-Copahue volcanic complex, Province of Neuquén, Argentina Johan C. Varekamp, J. Maarten deMoor, Matt D. Merrill, Anna S. Colvin, Adam R. Goss, Pieter Z. Vroon and David R. Hilton Geological Society of America Special Papers 2006;407;317-342 doi: 10.1130/2006.2407(15)
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Geological Society of America Special Paper 407 2006
Geochemistry and isotopic characteristics of the Caviahue-Copahue volcanic complex, Province of Neuquén, Argentina Johan C. Varekamp* J. Maarten deMoor Matt D. Merrill Anna S. Colvin Adam R. Goss* Earth and Environmental Sciences, 265 Church Street, Wesleyan University, Middletown Connecticut 06459-0139, USA Pieter Z. Vroon Institute for Earth and Life Sciences, Free University, DeBoelelaan 1085, Amsterdam 1081 HV, The Netherlands David R. Hilton Geological Research Division, Scripps Institution of Oceanography, La Jolla, California 92093-0244, USA
ABSTRACT The large Pliocene Caviahue caldera and associated active Copahue volcano are major volcanic features on the northwestern side of the Neuquén Basin. Chemical and petrographic data from the volcanic complex show a compositional range from basaltic andesite to rhyolite, predominance of two-pyroxene andesites and dacites, and a major quartz-biotite–bearing rhyolitic ignimbrite, which is part of the Riscos Bayos ignimbrite complex. Caldera wall sequences are dominated by lava and debris flows that show no consistent temporal trend toward more evolved magmas. The lavas at the top of the caldera wall series are among the most mafic in the region. The Copahue rocks are largely mafic two-pyroxene andesites, enriched in large ion lithophile elements and high field strength elements compared to the older Caviahue rocks. The intracaldera silicic rocks differ in composition from the Riscos Bayos ignimbrite sequence, which fills a paleovalley southeast of the caldera. The volume of these ignimbrites is insufficient to explain the formation of the Caviahue caldera. Cinder cones and lava flows east of the Caviahue complex consist of olivine-rich basalts. Isotopic data (Pb, Sr, Nd) show that all of the volcanic rocks in the Caviahue-Copahue volcanic complex have crustal components in their magma sources (i.e., subducted sediment and assimilated continental crust). The Caviahue series shows an increase in Sr and Pb isotopic ratios from mafic to silicic members, suggesting open-system evolution in the crust, although the isotopic variations are very small. The Copahue rocks are chemically and isotopically distinct from the Caviahue series; the He isotopic composition of geothermal gases from Copahue is close to mantle values, despite the evidence for sediment *E-mail: Varekamp—
[email protected]; present address, Goss— Department of Earth and Atmospheric Sciences, Snee Hall, Cornell University, Ithaca, New York 14853, USA.
Varekamp, J.C., Maarten deMoor, J., Merrill, M.D., Colvin, A.S., Goss, A.R., Vroon, P.Z., and Hilton, D.R., 2006, Geochemistry and isotopic characteristics of the Caviahue-Copahue volcanic complex, Province of Neuquén, Argentina, 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. 317–342, doi: 10.1130/2006.2407(15). For permission to copy, contact
[email protected]. ©2006 Geological Society of America. All rights reserved.
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J.C. Varekamp et al. involvement in the other radiogenic isotopic systems. The compositional differences between the Caviahue and Copahue series may be related to crustal processes, involving continental crust with a different chemical and isotopic composition for the two series. Alternatively, the change from Caviahue to Copahue volcanism might have been the result of a change in the element extraction process from the subducted complex: a contribution of sediment melting for Copahue versus volatile flux-induced mantle melting for the Caviahue series. The switch from a dominantly flux-melting regime (Caviahue series) to a drier decompressional melting regime led to smaller degrees of melting, as indicated by higher La/Yb values in the Copahue rocks. The low Ba/La values and high He isotope ratios in the Copahue series suggest an earlier phase of volatile fluxing and element loss from the subducted sediment complex, which probably took place below the main volcanic arc west of Copahue. As a result, the Copahue volcano east of the main volcanic front erupted magmas that formed in a drier mantle environment dominated by decompressional melting with a more significant component of subducted sediment melt. Keywords: Andes, volcanology, subduction, isotope geochemistry, andesites.
INTRODUCTION The active Volcán Copahue (37°51′S and 70°80′W) is located within the Andean Southern volcanic zone, approximately ~30 km east of the main volcanic front, as defined by the NNE-trending volcanic centers of Chillán, Antuco, Callaqui, Lonquimay, and Llaima (Fig. 1). It is located on the southwestern edge of an elliptic Pliocene caldera (19 × 15 km, Caviahue caldera), 30 km west of the town of Loncopué, Argentina (Fig. 2). This caldera is also referred to as the El Agrio Caldera or Caldera del Agrio in this volume and in the recent literature. The Copahue-Caviahue volcanic complex includes all the Pliocene to Holocene volcanic deposits associated with the formation of the caldera and volcano edifice construction. Unlike many stratovolcanoes in the Southern volcanic zone, Copahue has a unique concave-down shape, resembling a Hawaiian shield volcano. Delpino and Bermúdez (1993) attributed the volcano’s shape to glacial action during the Pleistocene, whereas Varekamp et al. (2001) suggested that repeated flank collapses, which resulted from progressive hydrothermal weakening of the interior, played a key role in the edifice morphology. We investigated the chemistry of the Caviahue and Copahue rocks, the origin of the Caviahue caldera, and how subducted components may have influenced this distal subduction environment behind the main arc. The combined processes of sediment subduction and crustal assimilation during magma genesis and evolution can be deciphered (e.g., Vroon et al., 1993), but this is more difficult with limited local sediment and crustal composition data. The rocks of the Caviahue caldera and Copahue volcano have been described previously (Bermúdez et al., 1993, 2002; Colvin, 2004; Delpino and Bermúdez, 1993; Bermúdez and Delpino, 2002; deMoor, 2003; Merrill, 2003; Goss, 2001; Linares et al., 1999), but comprehensive chemical and isotopic data have been lacking.
We present new major- and trace-element data and radiogenic isotopic data on the lavas of the Caviahue caldera walls, the Riscos Bayos ignimbrite sequence on the SE flanks of the Caviahue caldera slopes, and the intra- and extracaldera silicic rocks and the Copahue volcanic sequence. In addition, we present data for a backarc olivine-rich basaltic cinder cone from the Caviahue zone (Fig. 2). REGIONAL AND TECTONIC SETTING The Caviahue and Copahue volcanic rocks form part of the Southern volcanic zone of the Andes and are related to the eastward subduction of the Nazca plate below the South American continent. Fault-plane solutions indicate that between 0º and 49ºS, the Nazca plate is currently being subducted beneath South America in a roughly E-W compressional direction of 77 ± 12º at a rate of 9.2–10.8 cm/yr. There is little internal variation along strike in either direction and rate along the western South American continental margin (Pardo-Casas and Molnar, 1987). Volcanic gaps separate the Southern volcanic zone from the nonmagmatic Chilean flat-slab to the north and the Austral volcanic zone to the south (Jordan et al., 1983; Bevis and Isacks, 1984; Kay et al., 1987). The Southern volcanic zone is commonly divided into the southern Southern volcanic zone (from ~37°S southward), the transitional Southern volcanic zone (from 37°S northward), and the northern Southern volcanic zone (from 35°S north). The Mocha fracture zone collides with the trench at 37ºS and separates 36 Ma oceanic crust to the north from younger, 18 Ma crust to the south (Swift and Carr, 1974; Herron, 1981). Seismic profiles measured across the Southern volcanic zone indicate that earthquakes defining the Benioff-Wadati zone below the Southern volcanic zone initiate ~270–285 km east of the trench. Depths of intermediate-level earthquakes under Copahue volcano cluster between 90 and 120 km depth (Barazangi and Isacks, 1976;
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Figure 1. Regional map of the southern and transitional Southern volcanic zone (SSVZ and TSVZ, respectively) of the Andes, showing Copahue and the Caviahue caldera as the main volcanic front of active volcanoes.
Bevis and Isacks, 1984; Cahill and Isacks, 1992). The subducting slab has a relatively constant dip of 30º along the entire length of the Southern volcanic zone. The furthest eastward extent of intermediate-depth earthquakes in the Southern volcanic zone is directly beneath the modern volcanic arc. Backarc earthquakes are sparse in the Southern volcanic zone (Cahill and Isacks, 1992), which suggests that the slab is relatively warm below that region. The thickness of the overriding continental crust changes markedly at the 37ºS Mocha fracture zone boundary (Beck et al., 1996). South of 37ºS, the crustal thickness is relatively uniform at 35 km (Hildreth and Moorbath, 1988), but gravity data show a steady increase in crustal thickness from 35 km at 37ºS to 55 km at 34.5ºS (Dragicevic et al., 1961; Diez-Rodriguez and Introcaso, 1986). The age of the crustal rocks also increases to the north, providing a more radiogenic isotopic end member for magmas that have reacted with upper crust in this zone. Geochemical and isotopic signatures of crustal contamination are evident in magmas from volcanic edifices north of 37ºS (Hickey et al., 1986; Hildreth and Moorbath, 1988). Davidson et al. (1987, 1988) argued for a strong crustal influence south of 37° S as well. The Copahue-Caviahue volcanic complex marks the northern extent of the Copahue–Pino Hachado Precordilleran uplift. This uplift diverges to the southeast from the Andean Main
Cordillera south of 38º S. Its western limit is delineated by the Bío Bío–Aluminé fault system, which merges with the northernmost extension of the Liquiñe-Ofqui fault just south of the Copahue-Caviahue volcanic complex (Cembrano et al., 1996; Muñoz and Stern, 1988, 1989). The N-S–running Agrio valley bounds the uplift to the east and is marked by the NNW-SSE– trending Cordillera del Viento fault. Normal faults form the boundaries of the Pino Hachado uplift, and the low area to the east is called the “Graben of Loncopué” (Ramos, 1978), which has scoria cones and lava flows of backarc basalts. Quaternary volcanic centers are absent along the Cordillera del Viento fault at this latitude, yet ~150 km north along the fault are the massive Quaternary volcanic complexes of Tromen and Domuyo. Throughout the Pliocene and early Pleistocene, magmatism persisted within and along the margins of the Copahue– Pino Hachado uplift, as evidenced by NNW-trending centers between the Copahue-Caviahue volcanic complex in the north and Palao Mahuida in the south (Muñoz and Stern, 1989). A number of these Pliocene to Pleistocene centers have experienced caldera collapse, including the Copahue-Caviahue volcanic complex, Pino Hachado, and Palao Mahuida. In the early Pleistocene, volcanic centers on the southeastern margin of the uplift erupted alkali basaltic lavas, and by the late Pleistocene, volcanic activity had largely shifted to the west (Muñoz and
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Figure 2. Map of the Copahue-Caviahue volcanic complex showing the topographic rim of Caviahue caldera, Volcan Copahue, the Riscos Bayos ignimbrite sample localities, the Cerro Bayo dome, the sampled caldera sections, and the locations of the backarc cinder cones. Also shown are local drainages and the Las Maquinas geothermal area, which was sampled for gases.
Stern, 1989). Quaternary volcanism persists at the CopahueCaviahue volcanic complex and at fissure-aligned backarc cones in the valleys flanking the Copahue–Pino Hachado uplift (Moreno et al., 1986). REGIONAL VOLCANOLOGY Earlier volcanological studies (Pesce, 1989; Delpino and Bermúdez, 1993; Bermúdez and Delpino, 1999; Bermúdez et al., 2002; Linares et al., 1999; Múnoz and Stern, 1989; Mazzoni and Licitra, 2000) have named the Copahue-Caviahue volcanic complex as Volcán Copahue and the Copahue-Caviahue volcano and have drawn up schemes of evolution with local stratigraphic units. We use the following broad nomenclature: the Copahue-Caviahue volcanic complex consists of the older precaldera Caviahue sequences (e.g., Hualcupén Formation of Pesce, 1989), the intracaldera volcanic rocks, the Riscos Bayos ignimbrite sequence, the Cerro Bayo dome complex, the modern Copahue volcano, and the cinder cones and flows in the Loncopué graben area (Fig. 2). Pesce (1989) and Delpino and Bermúdez (1993) described four general eruptive episodes of volcanic activity at the Copahue-Caviahue volcanic complex labeled Copahue eruptive episodes I, II, III, and IV. The Copahue I phase began in the early Pliocene (4.3 ± 0.6 Ma), although Linares et al. (1999) also reported some ages
ca. 5 Ma, where the growth of numerous polygenetic stratovolcanoes was characterized by large-scale Plinian eruptions. Andesitic to rhyolitic pyroclastic and debris-flow deposits of this volcanic stage are widely distributed in both Argentine and Chilean territory. The Caviahue caldera formed during the Pliocene-Pleistocene, probably around 2.0 Ma (Linares et al., 1999; Folguera and Ramos, 2000). The stratigraphic record does not show large ash-flow deposits: the main exposed ignimbrite deposits on the Argentine side of the region, the Riscos Bayos flows near Loncopué, have probably <10% of the volume of the caldera. The caldera forms a major collapse structure in this region, and it formed the basement for the later evolutionary stages of Volcán Copahue (Muñoz and Stern, 1988). Rhyolitic domes that intruded into the floor of the caldera terminated this phase of magmatism (Niemeyer and Muñoz, 1986). Copahue II commenced during the Pleistocene (0.8 ± 0.1 Ma) with the nucleation of a new volcanic edifice on the southwest rim of the caldera. This new volcanic cone forms the underlying structure of the present-day Volcán Copahue. Deposits from this eruptive stage are found within the caldera in Argentine territory and outside of the caldera in Chile (Muñoz and Stern, 1988). During this eruptive episode, much of the volcanic cone was completely glaciated, leaving U-shaped valleys and striated lava flows. Evidence from striations on the exposed caldera walls and on intracaldera lava flows (on the peninsula) suggests that the
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Geochemistry and isotopic characteristics of the Caviahue-Copahue volcanic complex entire caldera could have been filled with ice during this period. Copahue III began just as the glacial cap began to recede. Lava flows and pyroclastic eruptions originated at the base of the still-glaciated Volcán Copahue and flowed into river valleys. Soon after, the glacier retreated to form the 100-m-thick glacier currently present at the summit of the volcano (Delpino and Bermúdez, 1993). Copahue IV includes the postglacial Holocene lava and pyroclastic deposits erupted largely from a NW-SE–trending, 2.5-km-long fissure near the summit of Volcán Copahue. Aligned along this fissure are several cones located inside the larger extinct crater. The youngest lava flows on the eastern flanks of the volcano belong to this stage and were erupted from a fissure on the north margin of the Río Agrio (Delpino and Bermúdez, 1993). The 2000 Copahue eruption produced mainly ashes and larger-sized (10–20 cm) clasts, which form a modern surface deposit mantling ice and older topography, with only minor preservation potential. In this study, we label the Copahue I volcanic rocks the “Caviahue series,” and the Copahue II, III, and IV, the “Copahue series.” Folguera and Ramos (2000) suggested that subsidence of the Caviahue caldera occurred as a tectonic manifestation of transtensional stress associated with strike-slip motion of the faults bounding the Copahue–Pino Hachado uplift. Pesce (1989), Delpino and Bermúdez (1993), and Linares et al. (1999) argued for a possible volcanic caldera collapse after the eruption of the Riscos Bayos ash flows.
321
2003; Colvin, 2004; Figs. 2, 3, and 4): (1) A dark red, finegrained basal layer with black pumices with quench rims, where the muddy groundmass was possibly wet during deposition (RB0). (2) A gray flow unit (RB1) made up of small (1–20 mm) fragments of white and gray pumice and black obsidian, with rare dark pumices in a matrix of ash and crystals. Hydrothermally altered lithics of volcanic origin occur as small clasts in this layer, which is up to 25 m thick. (3) A pale yellow (tan) flow unit (RB2) with floating mauve and dark pumices, up to 30 cm in size covering the gray flow, but part of the same cooling unit. Unaltered lithics of andesitic rocks several centimeters in size are common in this layer, which is up to 40 m thick. (4) The highest unit consists of bright white, highly indurated ash and crystals with small floating pink pumice clasts (RB3). This rock forms steep, columnar jointed cliffs of up to 20 m height,
FEATURES OF THE CAVIAHUE-COPAHUE VOLCANIC COMPLEX Field Relations and Petrography We sampled sections up the northern and eastern Caviahue caldera walls, the Riscos Bayos ignimbrites, the Cerro Bayo dome complex, and the intracaldera lavas on the peninsula that separates the two arms of Lake Caviahue (Fig. 2). Lavas from the three major eruptive episodes of Copahue volcano were sampled, including clasts from the 2000 eruption. A basaltic cinder cone with a dike was sampled where the Lower Río Agrio meets the Río Norquin, ~30 km north of Loncopué (Fig. 2). The northern Caviahue caldera wall is up to 600 m high and consists of a stack of lava flows interbedded with debris flows and thin pyroclastic deposits. The eastern wall has a sequence of lavas with thicker ignimbritic intercalations at the base, a group of trachyandesitic flows in the middle, and a top cover of mafic lavas. Many dikes cut the north and east wall sequences. About 40 samples were collected from lavas and ignimbrites in four vertical sections in these two caldera walls (deMoor, 2003; Merrill, 2003). The Riscos Bayos (RB) ignimbrites are mainly exposed in the Hualcupén canyon, ~20 km east of the southern exit of the caldera, and appear as a massive valley fill with several depositional units, listed here from topographically low to high (Merrill,
Figure 3. Valley fill with the Riscos Bayos ignimbrite units, with units RB0, RB1, and RB2 indicated. Note the continuous transition from RB1 to RB2.
Figure 4. Ridges of the RB3 ignimbrite, with a valley filled with the RB1 and RB2 sequence. The RB3 ridges may be former valley fills as well.
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J.C. Varekamp et al.
which form irregular E-W–trending ridges ~50–100 m wide. These rocks have been quarried for hundreds of years and gave rise to the local name Riscos Bayos (tan cliffs), where the “tan” refers to the color of a locally popular horse breed. The RB3 ridges and inselbergs occur far to the east and south on top of the stack of Caviahue lavas and debris flows, which are exposed in the Hualcupén canyon walls (Fig. 2). The stratigraphic relationship between the Riscos Bayos units is complex, but it appears that valleys in between the RB3 ridges contain topographically lower outcrops of RB1 and RB2 units (Fig. 4). The RB3 unit, although topographically the highest, is most likely the oldest Riscos Bayos flow, and valleys cut into either a flat RB3 ignimbrite plateau, or E-W valley fills of RB3, now present as E-W ridges after reversal of topography, were then later filled with the RB0, RB1, and RB2 flows. The maximal estimated total dense rock volume of the Riscos Bayos ignimbrites (~7 km3) is much less than that of the caldera (160 km3), and the Riscos Bayos eruptions cannot have been the main cause of the caldera collapse. The Cerro Bayo dome on the north slope of the Caviahue caldera sequence (Fig. 2) postdates the formation of the caldera. The peninsula between the two arms of Lake Caviahue inside the caldera has a dense, black, and glassy flow at the bottom, which becomes more vesicular toward the top. The flow morphology suggests that the peninsula flow filled a former valley between the areas where we now find the two 100-m-deep arms of Lake Caviahue, suggesting a radical inversion of relief over time. The top of the peninsula sequence and the many loose boulders and ridges created by glacial processes consist of black streaked material that resembles a welded ignimbrite (Mazzoni and Licitra, 2000). These deposits could also be interpreted as deposits from a gassy lava flow or from lava fountaining (Mazzoni and Licitra, 2000; Bachmann et al., 2000), where black, highly degassed blobs of magma (now obsidian) were carried as pseudoliquid clasts. These clasts were then slightly deformed during transport and deposition, which means they are not the fiamme of welded ignimbrites. A large fan of steeply bedded siliciclastic materials make up the S-rim of Lake Caviahue and also crop out on the peninsula; this may be a hydromagmatic surge facies of the peninsula sequence. The lava-flow deposits are found on both sides of the southern arm of Lake Caviahue: either the lake was not there yet at the time of deposition, or it was filled with ice, and the flows covered the ice, leading to the hydromagmatic deposits. The peninsula lavas and hydroclastic deposits are an intracaldera sequence, which postdates the collapse. Small ignimbrite outcrops occur along the west wall of the caldera, but most of these rocks have been altered into yellow clay and cannot be correlated with any other field units. Much of the basic geology and geomorphology of the Caviahue caldera is poorly documented, e.g., the age relations between the caldera collapse, the formation of the two canyon outlets to the east, and the formation of the valleys that now contain Lake Caviahue are as yet undetermined. The south and west walls of
the Caviahue caldera are less well defined, consist of large isolated cliffs and walls, as described by Todd (2005). The broad base of Copahue volcano is made of Copahue III mafic lava flows. On top, a sequence of striated lava flows is found (Copahue II), including those that form a series of cascades of the upper Río Agrio (Fig. 5). The modern, postglacial flows and ashes of Copahue form the top unit, which includes the 2000 eruptive products. Cinder cones with dikes and lava flows of olivine-rich basalt occur east of the Copahue-Caviahue volcanic complex (Bermúdez et al., 1993; deMoor, 2003), and these units postdate the Caviahue caldera formation (Múnoz and Stern, 1989). Age relations between the field units are known from K-Ar dating: The caldera lava sequence has ages between 5 and 2 Ma, and samples from the Riscos Bayos RB3 unit have been dated at 2.05 Ma (Múnoz and Stern, 1988; Linares et al., 1999). The Cerro Bayo rocks have an age of ca. 0.62 Ma (Linares et al., 1999). The Copahue activity probably started around 0.8 Ma (Delpino and Bermúdez, 1993). The lavas of the Caviahue caldera walls range from olivine-rich basaltic andesites to two-pyroxene andesites and dacites. Some rare lava flows carry only plagioclase and clinopyroxene, and some have an abundance of plagioclase crystals only. The siliceous ignimbrites carry mainly plagioclase, two pyroxenes, opaques, and very rarely biotite (one sample). Phenocryst compositions have been reported by deMoor (2003) and Merrill (2003). The RB1–RB2 samples are two-pyroxene andesitic pumices, whereas the RB3 flow has abundant embayed quartz and pristine biotite crystals. The intracaldera peninsula rocks are rich in plagioclase with orthopyroxene, less common clinopyroxene, and rare olivine. The Copahue sequence has olivine-bearing and two-pyroxene andesites, but no evolved rocks or samples with hydrous pheno-
Figure 5. Cascades in the upper Río Agrio valley in columnar jointed lava flows from Copahue eruptive episode II, just west of Caviahue village.
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Geochemistry and isotopic characteristics of the Caviahue-Copahue volcanic complex crysts were encountered (Goss, 2001). The backarc cinder cones carry abundant olivine and clinopyroxene with basic plagioclase and opaques. Detailed accounts of the petrography, phenocryst compositions, and glass inclusion studies can be found in deMoor (2003), Merrill (2003), and Goss (2001). Major-Element Rock Chemistry and Classification Whole-rock analyses of Copahue-Caviahue volcanic complex samples (Tables 1–4) show a large compositional range of the volcanic rocks in this region (Figs. 6A and 6B). The lavas and ash flows of the caldera wall sections vary from basaltic andesite to rhyolites, where sample EB-8 is the least-evolved east caldera wall rock, with 4.9% MgO, and sample NC8.5, a basaltic andesite lava flow rich in olivine in the north caldera wall, has 7.4% MgO. Most intermediate rocks from the caldera walls are trachyandesitic with more-evolved trachydacites and rhyolites. The three main units of the Riscos Bayos ignimbrites are respectively trachydacites (RB1), trachyandesite (RB2), and rhyolites (for the quartz-bearing RB3 unit). Some individual clasts in RB1 (dark scoria–pumice, RB1–3a) have a composition similar to the RB2 pumice sample. The silicic RB3 unit has no inclusions of other stratigraphic units. Some of the ignimbrite layers that outcrop low in the caldera walls are rhyolites similar to RB3 in bulk rock composition, and one unit also carries biotite crystals (deMoor, 2003). The intracaldera rocks from the peninsula between the two arms of Lake Caviahue are largely trachydacites that are slightly more evolved than the RB1 unit; our samples from the peninsula show substantially more compositional variation than the samples analyzed by Mazzoni and Licitra (2000). The Cerro Bayo dome rocks have silica concentrations comparable to RB3 (72% SiO2) but lack the hydrous phenocrysts and quartz crystals. The Copahue rocks range from 55% to 62% SiO2 and have MgO concentrations in the lowest stratigraphical units of ~4.1% and between 4.1% and 4.4% for scoria from the 2000 eruptions. All other Copahue rocks are slightly more evolved, with MgO from 2.4% to 3.9%. The K2O contents of Copahue rocks are relatively high, from 1.7% to 3.0%. The mafic lavas and scoria with ~50% SiO2 and 5.6%–6.4% MgO from the backarc classify as basalts (Fig. 6A). Overall, the Copahue-Caviahue volcanic complex rocks are alkali-rich, and most fall in the medium-K calc-alkaline field, whereas most of the Copahue and evolved Caviahue rocks plot in the high-K calc-alkaline field (Fig. 6B). The CopahueCaviahue volcanic complex products are enriched in alkalis compared to the southern Southern volcanic zone and plot largely with the transitional Southern volcanic zone rocks (in agreement with their location on the boundary of the transitional Southern volcanic zone and southern Southern volcanic zone [37°S]). No shoshonitic or other exotic lavas were encountered in this study, despite the position of the Copahue-Caviahue volcanic complex at the far side of the arc (Gill, 1981). Major-element variation diagrams show the chemical systematics of the Copahue-Caviahue volcanic complex rocks
323
with respect to each other and within the regional context. Most MgO versus major-element plots (Fig. 7A–C) show a relatively smooth trend for the Caviahue rocks, including the Riscos Bayos, backarc basalts, and intracaldera lavas. The RB3 unit always plots off this trend. The Copahue rocks show separate trends outside the Caviahue caldera array. Initial least square crystal fractionation modeling suggests that many of the Copahue rocks can be related to each other by crystal fractionation (Goss, 2001). Not enough microprobe data are available for the Caviahue caldera rocks to make this point, but the coherent trends and similar mineralogy of the rocks suggest close ties. The quartz-bearing RB3 rhyolite (Merrill, 2003; Colvin, 2004) is the most-evolved rock and plots off the general array trend. The intracaldera samples from this study plot separately from the Riscos Bayos group, and, based on majorelement data, are not the intracaldera equivalents of the Riscos Bayos units, contrary to arguments by Mazzoni and Licitra (2000). The Cerro Bayo sample plots close to the RB3 samples, but has different mineralogy. The Caviahue rocks have the greatest similarity to the southern Southern volcanic zone rocks (low K2O, high CaO, Fe2O3, and TiO2), whereas the Copahue rocks plot within the fields of the transitional and northern Southern volcanic zone (high K2O, low CaO, Fe2O3, TiO2). These trends are significant over the whole compositional range of each family. Trace-Element Chemistry The trace-element chemistry of the Copahue-Caviahue volcanic complex rocks is given in Tables 1–4 and is illustrated in MgO versus trace-element diagrams (Fig. 8), and in relative abundance diagrams for incompatible elements and rare earth elements (REEs) (Fig. 9). As with the major elements, a fundamental division between Copahue and Caviahue rocks is obvious from the MgO versus trace-element diagrams. The elements Ba, Cs, La, Nb, Rb, U, Th, and Zr are enriched in the Copahue series relative to the Caviahue rocks at a given MgO level, whereas similar concentrations of Sr and Yb are found in both groups. The enrichment in highly incompatible elements in the Copahue series agrees with the more potassic nature of these magmas. The Cerro Bayo sample (labeled as an IC sample in these figures) is very similar to the RB3 unit, whereas the intracaldera peninsula rocks are distinct from the RB1–RB2 units. In all diagrams both rock suites have a gentle trend of increasing incompatible element concentrations with decreasing MgO contents, whereas the RB3 (both bulk rock and single pumice samples) and Cerro Bayo rocks have much higher concentrations, well outside that trend. The trace-element ratio Nb/Yb is constant in much of the southern Southern volcanic zone volcanic rocks (Fig. 8F), and many of the samples from the Caviahue series plot in that same field. The Riscos Bayos and backarc basalts samples hover just above this field, with the RB3 units outside the array as usual. The Copahue rocks have much higher Nb/Yb ratios at similar
24.5
2.6
Y
Yb
a
a
164
3.7
34.5
1.6
5.6
451
39.4
10
30.3
5
20.9
1.7
13
49
455
0.3
1.2
8.3
1.2
0.3
1.1
8.6
1.4
2.0
6.0
1.5
8.6
4.8
2.3
0.5
3.4
4.2
1.3
7.2
4.2
1.5
0.2
1.0
8.8
1.4
9.3
2.9
7.4
a
a
a
a
a
5.7
7
75
1.5
38
30
5
10
4
22.7 33.6 13.1
0.9
17
51
10
18
7
6 3.7
13.1 1.1
4.6
4.3
4
2.1
38.4 35.6 20.1
1.7
a
a
a
a
128 166 299 121
2.2
22
0.9
3.5
601 385 215 396
20.9 40.2 79.6 19.1
8
22.1 31.5 41.7 18.5
4
15
1.3
35
36
302 441 490 237
a
a
100.1
0.3
1.6
8.3
1.2
9.2
3.7
4.8
18.4
52.9
a
226
4.4
40.1
2.5
8.5
341
59.6
17
40.3
7
30.5
2.5
5
60
555
a
98.6
0.4
2.4
3.8
1.2
6.3
5.0
1.3
16.7
61.9
a
191
2.8
25
2.2
7.6
429
50.1
15
27.4
6
22
3.4
23
49
417
a
98.8
2.0
6.6
1.3
7.8
3.3
4.1
17.4
56.3
c
65
3.5
35.8
1.8
7.6
819
5
13
37.2
8
26.9
3
19
59
247
b
100.3
0.2
1.7
6.9
1.6
9.7
3.9
3.1
15.9
56.1
c
241
4.9
48.8
2.8
10.5
310
52
17
52.3
13
39.3
3.8
7
84
576
b
100.0
0.5
2.8
4.3
1.5
7.0
4.4
1.8
15.5
61.9
2.2
4.9
1.2
7.5
5.1
2.0
b
113
2.2
21.8
0.8
3.5
535
23
7
21.1
4
15.3
1.3
29
32
291
c
d
99.9 99.5
0.3
1.0
8.7
1.1
9.3
3.5
4.9
18.3 17.3
52.8 59.3
413 d
5.5
49
3.7
15.5
255
111.5
24
71.7
15.5
64.3
4.2
3
104
843
d
99.6
3.6
1.9
0.7
3.4
6.3
0.5
16.5
66.7
d
99.5
2.0
5.0
0.8
6.2
5.3
2.1
17.6
60.6
d
100.2
4.8
1.3
0.3
1.6
4.0
0.2
14.1
74.0
CPEN0
10.4
41.4
17
42.4
9
6
84
719
c
184
3.5
30
1.6
8.2
543
b
99.7
0.1
3.9
2.6
0.8
5.3
4.7
1.0
15.2
65.7
420 a
4.2
38
5
0.6
219
26.2 121
10
26.4
6
20.7
1.5
17
46
365
c
99.6
0.3
1.2
7.9
1.5
7.6
3.8
2.8
18.9
52.1
418 a
4.5
42.2
5.2
0.7
227
108
12
47.2
15
49.1
2.6
6
92
747
b
99.7
0.1
3.7
2.5
0.8
5.2
4.7
0.9
15.4
66.0
CPEN1-2
a
414
5.3
50.5
5
0.8
240
114
13.2
52.4
14
54.7
5.7
5
100
817
b
99.8
0.1
3.8
2.3
0.8
5.1
4.8
0.8
15.5
65.7
KCP2
297 a
3.4
30.2
3.4
0.5
395
72
8.7
35.2
12
35.9
14.9
14
70
665
b
100.1
0.1
2.4
4.5
1.2
7.3
3.9
2.2
17.8
60.0
CPEN3B
Notes: Rock samples were screened based on loss on ignition (LOI) contents and petrographic study. Selected samples (38) were analyzed for major elements by wavelength-dispersive X-ray fluorescence (XRF, Norelco instrument) at Wesleyan University, using a set of international standards for calibration (AGV, BHVO, BCR). Most samples were analyzed for trace elements by ICP-MS by a contract laboratory (SGS, Toronto, Canada). Sources: a—deMoor (2003) b—Merrill (2003); c—Colvin (2004); d—Goss (2001).
Source
144
1.8
U
Zr
6.4
Th
562
39.9
Rb
Sr
8
Pb
18.9
La
5
1
Cs
25.1
25
Co
Nd
43
Ce
Nb
386
Ba
a
Source
3.2
5.5
9.3 10.1
3.5
5.5
99.0 99.1 99.7 98.8 98.5 99.4
0.3
1.1
8.7
1.3
10.1
3.3
5.9
17.7 18.6 18.3 16.6 16.0 16.6
51.0 51.6 51.4 57.1 60.8 52.0
Peninsula Lavas
324
a
99.0
1.8
98.8
1.5
K2O
5.4
Total
7.4
CaO
1.3
0.5
1.4
TiO2
7.4
4.8
1.9
17.6
58.8
NA6-1 NB2 NB3 NB4 NB5 NB6 NC8.5 NC9-1 NC11-2 NC12-1 EA3-1 EB2-1 EB-8 1-011299 2-011299 3-011299 1-021299 HCD
P2O5
3.8
9.5
3.2
MgO
Fe2O3
17.7
Al2O3
Na2O
54.2
NA4-1
SiO2
East Caldera Wall
TABLE 1. MAJOR- AND TRACE-ELEMENT COMPOSITIONS OF THE CAVIAHUE ROCK SERIES North Caldera Wall
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Geochemistry and isotopic characteristics of the Caviahue-Copahue volcanic complex
325
TABLE 2. WHOLE-ROCK COMPOSITIONS FOR MAJOR AND TRACE ELEMENTS IN PUMICE SAMPLES OF THE RISCOS BAYOS FORMATION RB1 RB1-1
RB1-2
RB1-3
RB2 RB1-3a
RB1-4
SiO2 Al2O3 MgO Na2O Fe2O3 TiO2 CaO K2O Total LOI Source
63.1 15.7 1.1 6.0 4.5 1.0 2.7 2.5 101.1 4.6 b
64.2 16.0 0.9 6.3 4.3 0.8 2.4 2.7 101.4 3.7 b
64.2 16.1 1.1 6.8 4.6 0.8 2.5 2.6 101.2 2.5 b
57.9 16.3 2.6 5.3 7.8 1.6 5.8 1.8 100.2 1.2 b
64.5 15.9 1.0 6.2 4.9 0.9 2.7 2.7 101.0 2.2 b
Cs Rb Ba Th U Nb Sr Hf Zr Y Pb Ta La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
3.9 62 552 7.4 2.2 10 293 7 285 38 10 0.8 31.1 69.5 9.1 38.2 8.2 2.2 8.1 1.2 7.4 1.5 4.2 0.63 4.3 0.67
4 63 564 8.1 2.3 10 269 8 299 38 17 0.8 32.5 72.6 9.4 39.2 8.2 2.1 8.0 1.3 7.2 1.5 4.5 0.65 4.6 0.7
3.9 62 556 7.7 2.2 10 278 8 299 38 10 0.8 31.4 70.8 9.1 38.6 8.1 2.2 8.3 1.2 7.4 1.5 4.4 0.62 4.4 0.68
2.4 41 490 6.1 1.6 8 472 5 185 37 bd 0.7 27.1 60.9 8.9 38 8.4 2.2 8.0 1.1 6.5 1.5 3.9 0.58 3.4 0.5
4.9 73 568 9.1 2.7 10 238 9 335 37 10 0.8 32.7 73.1 9.4 39 8.4 2.0 8.2 1.3 7.3 1.5 4.4 0.62 4.5 0.68
RB2-1 55.5 16.3 2.1 5.5 7.4 1.1 5.3 1.5 99.8 0.1 c
RB2-2 54.9 15.6 2.4 4.7 6.8 1.4 5.1 1.1 100.5 8.4 b
0.5 0.7 22 21 559 541 6.9 5.4 1.7 1.5 9 7 564 399 6 5 240 161 40 31 10 bd 0.7 0.6 32.1 23.6 74.4 53.1 10.0 7.4 43.1 32.8 9.2 7.2 2.7 1.9 9.4 6.5 1.4 1.0 8.1 5.6 1.6 1.3 4.5 3.7 0.63 0.52 4.2 3 0.67 0.43
RB3 RB-2-3
RB3-1
RB3-2
TR3-2
55.0 16.4 1.9 2.2 6.1 1.2 3.5 1.9 100.2 11.7 c
74.6 14.1 0.2 4.3 1.8 0.3 1.1 4.2 101.3 0.6 b
70.5 14.3 0.2 4.0 3.7 0.3 1.3 3.6 98.9 1.0 b
74.1 14.1 0.1 5.2 1.5 0.2 0.9 4.4 101.4 1.0 b
2.8 45 812 7.3 1.6 9 459 6 210 35 bd 0.7 28.1 61.3 9.4 40.2 8.8 2.2 7.7 1.2 7.8 1.7 5.1 0.75 4.7 0.68
7.1 100 546 15.6 4.5 7 99 3 119 12 11 0.9 24.3 41.9 4.6 14.7 2.4 0.4 2.3 0.4 2.1 0.5 1.4 0.21 1.6 0.23
4.9 105 550 18.1 5.5 14 127 4 133 22 16 1.2 30 50 5.7 19.6 3.6 0.6 3.9 0.7 4.0 0.9 2.6 0.41 2.7 0.43
4.8 93 670 23.7 5.4 9 92 5 137 15 16 1.2 28.6 48.2 5.1 16.9 2.9 0.5 2.4 0.4 2.2 0.4 1.4 0.22 1.6 0.24
Note: LOI—loss on ignition, Samples RB3-1 and TR3-2 are bulk ignimbrite samples, sample RB3-2 consists of individual pumice clasts. Sources: b—Merrill, 2003; c—Colvin, 2004; bd—below detection.
MgO contents and plot in the field of the transitional Southern volcanic zone and possibly the northern Southern volcanic zone. Relative abundance diagrams for the local rock units (Fig. 9A) show very similar patterns, with strong negative anomalies for Ti, P, and Nb-Ta. This is true from the basaltic backarc basalt samples to the rhyolites of RB3. The Cerro Bayo sample (IC) and RB3 both show anomalously strong enrichment in U and Th (both bulk and pumice RB3 samples), whereas the backarc basalts and the mafic Caviahue caldera lava show the lowest enrichments in the most incompatible elements. The RB3 samples are relatively depleted in the more compatible elements of this suite of trace elements. The element Sr shows various degrees of depletion, with a small positive spike for the backarc basalt sample. The REE diagrams show strong family ties between all Copahue-Caviahue volcanic complex samples (Fig. 9B), with
gently sloping heavy (H) REE curves, small negative Eu anomalies, and steeper light (L) REE patterns. The mafic Caviahue lava and the backarc basalt sample show the lowest LREE slopes. The RB1–RB2 series has the same pattern as the rest of the Caviahue and Copahue rocks, whereas the RB3 sample has a flat HREE slope, much steeper LREE slope, and the deepest Eu troughs (Fig. 9C). The RB1 and RB2 samples are all very similar, but the three samples from the RB3 unit are different (Fig. 9C). These diagrams confirm the close family ties between the Copahue-Caviahue volcanic complex samples with the exception of RB3, which has the very steep LREE patterns and is chemically most similar to the postcaldera Cerro Bayo dome rocks (Fig. 9C). Detailed discussions of the REE chemistry are given by Colvin (2004), deMoor (2003), Merrill (2003), and Goss (2001).
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J.C. Varekamp et al. TABLE 3. WHOLE-ROCK ANALYSES FOR MAJOR AND TRACE ELEMENTS OF CERRO BAYO AND COPAHUE SAMPLES Cerro Bayo NE2-2
1-251199
2-251199
3-251199
56.1 17.1 3.9 3.8 7.9 1.2 7.2 2.2 99.4 d*
55.6 17.3 4.1 3.4 8.2 1.1 7.3 2.0 99.0 d*
SiO2 Al2O3 MgO Na2O Fe2O3 TiO2 CaO K2O Total Source
72.5 15.2 0.4 4.2 2.1 0.3 2.0 3.9 100.6 a
55.7 17.7 4.3 4.0 8.3 1.2 7.5 2.1 100.9 d
Ba Ce Co Cs La Nb Nd Pb Rb Sr Th U Y Yb Zr Ref
577 33.7 2.9 6.2 17.6 4 13 21 90 152 16 4.5 11 1.5 131 a
491 62.7 24.5 4.5 43.7 4 37 22 59 538 9 2.3 28 3.4 180 c
CEE II 4-251199 55.5 17.3 4.1 3.6 8.0 1.1 7.2 2.1 98.9 d
1-281199 57.6 17.7 4.1 3.4 8.1 1.2 7.2 2.0 101.4 d*
3-281199 57.6 16.4 4.1 4.0 8.0 1.2 7.1 2.0 100.3 d*
55.4 17.4 3.9 3.4 8.2 1.2 7.9 1.9 99.3 d*
Lower CEE III 1-271199
2-271199
56.8 16.8 3.6 4.9 7.9 1.3 6.4 2.3 99.9 d*
56.4 17.0 3.7 3.9 8.2 1.3 6.6 2.0 99.0 d*
c
c
19
c
9
12
30
33
c
c
c
Upper CEE III
c
c
Copahue E.E. IV
8-261199
5-271199
6-271199
7-271199
8-271199
1-291199
SiO2 Al2O3 MgO Na2O Fe2O3 TiO2 CaO K2O Total Source
58.4 17.6 3.1 4.4 7.0 1.1 6.2 2.5 100.2 d
59.5 16.9 2.8 4.5 7.0 1.2 5.2 2.8 99.9 d*
59.1 17.0 2.8 4.9 7.2 1.2 5.3 2.9 100.4 d*
62.2 17.0 2.4 3.8 7.2 1.2 4.6 3.0 101.4 d
54.9 17.6 3.7 4.2 8.8 1.2 6.9 1.7 98.9 d
59.1 17.1 2.5 4.9 7.0 1.1 5.5 2.7 99.8 d*
55.1 17.5 4.2 4.0 8.5 1.2 6.8 2.0 99.3 d
Ba Ce Co Cs La Nb Nd Pb Rb Sr Th U Y Yb Zr Ref
546 80.0 16.3 3.0 37.6 12 44 20 94 545 12 2.8 32 3.3 253 c
c
633 56.5 14.1 3.8 27.3 15 32 27 116 411 14 2.9 25 2.7 315 c
441 61.3 24.8 1.6 29.6 12 38 18 48 545 9 1.6 29 3.3 200 c
c
485 62.7 27.7 4.3 30.7 10 35 20 73 547 9 2.2 28 3.0 220 c
c
5-261199
1-261199
3-261199
2000 eruption COP1
COP3
COP7
55.4 17.6 4.1 3.6 8.3 1.2 7.1 2.0 99.3 d*
55.2 17.2 4.1 3.9 9.0 1.2 7.3 2.0 99.8 d
55.3 17.1 4.3 4.1 8.7 1.2 7.2 2.0 100.0 d
55.1 17.2 4.4 3.9 8.5 1.2 7.4 1.9 99.7 d*
c
446 59.8 26.9 3.5 28.4 10 31 14 60 488 11 2.6 28 3.0 194 c
446 59.9 27.2 3.4 27.7 10 32 13 61 485 12 2.6 28 2.9 205 c
c
Note: CEE—Copahue eruptive episode. Sources: a—deMoor ( 2003); d—Goss (2001). Some samples (indicated with *) were analyzed for major elements and for a limited set of trace elements by XRF by a contract laboratory (SGS, Toronto, Canada).
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Geochemistry and isotopic characteristics of the Caviahue-Copahue volcanic complex TABLE 4. WHOLE-ROCK ANALYSES OF BACK ARC BASALTS FOR MAJOR AND TRACE ELEMENTS AE3D AE3S AE3D AE3S SiO2 50.2 50.4 Ba 446 238 Al2O3 18.0 18.5 Ce 28 28 MgO 6.4 5.6 Co 32 30 Na2O 3.0 3.2 Cs 0.2 0.3 Fe2O3 10.0 9.9 La 13.5 13.1 TiO2 1.1 1.2 Nb 8 5 CaO 10.0 10.2 Nd 16.4 16 K2O 0.9 0.9 Pb 3.7 P2O5 0.2 0.3 Rb 46 11 Total 100.2 100.0 Sr 413 791 Th 0.2 1.7 U 0.6 0.5 Y 16.0 16.1 Yb 1.7 1.6 Zr 201 65 Source c a a a Note: Sources: c—Colvin (2004); a—deMoor (2003).
Isotope Geochemistry The radiogenic isotopic ratios of Sr, Nd, and Pb of the Copahue-Caviahue volcanic complex rocks (Table 5) provide information on the relations between the different CopahueCaviahue volcanic complex rock suites as well as their source components. A plot of 87Sr/ 86Sr versus latitude (Fig. 10) shows the regional relationship of high-Sr isotopic values in the northern Southern volcanic zone, lower values in the transitional Southern volcanic zone, and then low values farther south in the southern Southern volcanic zone. The new Copahue-Caviahue volcanic complex data plot very well within this regional context, with the backarc basalt sample at the lowest recorded 87Sr/ 86Sr in the Southern volcanic zone, slightly higher values for the Caviahue samples, a still higher value for the Copahue sample, and sample RB3 with the highest Sr isotope ratio. Our Sr isotope measurement for the backarc basalt sample is significantly lower than the values obtained by Múnoz and Stern (1989) on backarc lavas somewhat farther south (Laguna Blanca area), which also differ substantially in trace-element concentrations. The Sr-Nd isotope diagram (Fig. 11) shows the fields of mid-ocean-ridge basalt (MORB) and Nazca plate basalts, and that of detrital sediments from the Pacific Ocean (sources for all data and end members are given in the caption of Fig. 6). The northern, transitional, and southern Southern volcanic zone fields are indicated, as well as the Copahue-Caviahue volcanic complex samples analyzed for this study. The three Southern volcanic zone subzones show very tight groupings in this diagram, with the southern Southern volcanic zone samples closest to the MORB field and the northern Southern volcanic zone farthest away. The backarc basalt plots closest to the MORB end member, whereas the Caviahue samples (EB8, RB1, RB2, and RB3) plot toward the Pacific Ocean detrital sediment field. The Caviahue samples plot together with those from the southern Southern volcanic zone, whereas the Copahue sample plots in the transitional Southern volcanic zone array and off the
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Caviahue internal trend. The isotopic differences between the analyzed Caviahue samples are very small (except for sample RB3), but the differences between the Copahue, Caviahue, and backarc basalt samples are significant. As a first-order conclusion, all these arc rocks contain a crustal component, either from subducted materials, from erosion of the overriding plate, or from interaction with the local continental crust. The RB3 magma and Copahue rocks carry the largest crustal contributions, the other Caviahue samples have smaller amounts, and the backarc basalt has only a very small component. The nature of this crustal contaminant cannot be determined from this diagram alone, although Pacific Ocean detrital sediment seems a likely candidate. The 207Pb/ 204Pb versus 206Pb/ 204Pb diagram (Fig. 12A) shows the fields of MORB with the Northern Hemispheric reference line (NHRL), the Nazca plate basalts (NPB), and the Pacific Ocean detrital sediment field. Data from the northern, transitional, and southern Southern volcanic zone are plotted with the Copahue-Caviahue volcanic complex samples, but the three Southern volcanic zone subzones do not cluster in the Pb isotope diagrams as in the Sr-Nd isotope diagram, possibly as a result of variations in Pb isotopic composition of the subducted sediment. All Copahue-Caviahue volcanic complex samples plot within the Pacific Ocean detrital sediment field, indicating that the sedimentary or crustal Pb contribution overwhelmed the mantle isotopic Pb signature. A detail of the 207Pb/ 204Pb versus 206Pb/ 204Pb diagram (Fig. 12B) shows the relations between the Copahue-Caviahue volcanic complex units, with the Caviahue samples showing only very small variations in 207Pb/ 204Pb and 206Pb/ 204Pb, but the Copahue sample is well outside that range, apparently enriched in 206Pb. The Pb isotopic composition of the Copahue sample is not the result of a larger addition of a similar component that places the Caviahue samples away from the MORB field; the Copahue sample is predominantly displaced toward a higher 206Pb/ 204Pb value. The Copahue rock is influenced by a different source component compared to the crustal contaminant of the Caviahue samples. The 208Pb/ 204Pb versus 207Pb/ 204Pb diagram (Fig. 13A) shows once more that the Copahue-Caviahue volcanic complex and Southern volcanic zone samples are all displaced from the MORB or Nazca plate basalt (NPB) fields toward the Pacific Ocean detrital sediment field. The Caviahue samples plot just below the Pacific Ocean detrital sediment field in a cluster with other Southern volcanic zone samples. The Caviahue samples form a linear array (barely outside the analytical error), with RB3 at the most radiogenic end and the mafic lava EB8 at the least radiogenic end (Fig. 13B). We interpret this as evidence for crustal assimilation in the Caviahue series during crustal residence of the magmas, with the largest crustal contribution in the rhyolitic RB3 sample. The RB3 rocks are also strongly enriched in U and Th, and we hypothesize that the magma assimilated crustal material rich in U and Th that also carried the radiogenic daughters of these elements. The Copahue sample is displaced toward a 208Pb/ 204Pb–rich component just into the Pacific Ocean detrital sediment field, which rep-
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A
B
Figure 6. Rock classification diagrams for the Copahue-Caviahue volcanic complex samples (Irvine and Barangar, 1971). Note the broad range from basalts to rhyolite (A) and the slightly alkaline character of most rocks (B; LeBas et al., 1986). NSVZ—northern Southern volcanic zone; TSVZ—transitional Southern volcanic zone; SSVZ—southern Southern volcanic zone, RB—Riscos Bayos ignimbrite units; BAB—backarc basalt; IC—intracaldera samples. Data sources for these and all following diagrams: Davidson et al. (1988); Deruelle et al. (1983); El-Hinnawi et al. (1969); Frey et al. (1984); Futa and Stern (1988); Gerlach et al. (1988); Goldstein and O’Nions (1981); Harmon et al. (1984); Hickey et al. (1986); Hickey-Vargas (1989, 1995); Hildreth and Moorbath (1988); Jahn et al. (1980); Kilian and Behrmann (2003); Klerkx et al. (1977); Lopez-Escobar et al. (1977, 1991, 1995); Othman et al. (1989); Pichler and Zeil (1972); Plank and Langmuir (1993, 1998); Stern et al. (1990); Sun and McDonough (1989); Thompson and Bryan (1976); Thorpe and Francis (1979); Thorpe et al. (1982); Tormey et al. (1991); Unruh and Tatsumoto (1976); Winter (2001); and Worner et al. (1992).
resents a different process than the crustal assimilation that created the differences between the individual Caviahue samples. The Copahue samples are enriched in Th relative to the Caviahue series, as well as in 208Pb, the radiogenic daughter of Th. The two Pb isotope measurements from Pb dissolved in the Copahue hot-spring waters presumably represent an average of the Pb isotope signature of the whole Copahue rock series (Varekamp et al., 2001), and they have a very similar Pb isotopic signature to the 2000 Copahue rock sample analyzed in this study (expressed as 207Pb/ 206Pb and 208Pb/ 206Pb ratios, Fig. 14). The Copahue samples plot well away from the Caviahue samples, with an enrichment in 208Pb and 206Pb relative to the four
Caviahue samples. In this Pb isotope space, the Copahue sample carries a source component that, according to Pb-Pb ages, is younger than the component in the Caviahue samples. The conclusion from these Pb-Pb isotope diagrams is that the Caviahue magmas obtained enough Pb with the daughters of U (207 and 206) to make it resemble the isotopic composition of the Pacific Ocean detrital sediments, but not enough Pb of Th heritage (208) to plot inside the Pacific Ocean detrital sediment field. The Copahue sample carried a component more enriched in 207Pb and 206Pb, especially the latter, but also acquired enough of a 208-rich Pb component to plot just inside the Pacific Ocean detrital sediment field.
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Geochemistry and isotopic characteristics of the Caviahue-Copahue volcanic complex
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A
B
C
Figure 7. Major-element variation diagrams (MgO versus K2O [A], CaO [B], and Fe2O3 [C]) showing the coherent arrays for the mafic and intermediate Caviahue rocks, the more enriched RB3 and Cerro Bayo samples, and the overall similarity between the lower Riscos Bayos units and the caldera lava samples. The Copahue rocks stand out through enrichment in K2O and depletion in CaO and Fe2O3. The Copahue rocks tend to plot with northern Southern volcanic zone (NSVZ) and transitional Southern volcanic zone (TVSZ) samples, whereas the Caviahue rocks are like southern Southern volcanic zone (SSVZ) samples. RB— Riscos Bayos ignimbrite units; BAB—backarc basalt; IC—intracaldera samples.
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A
B
C
Figure 8. Trace-element variation diagrams (MgO versus Ba [A], Rb [B], Nb [C], Th [D], Sr [E], and Nb/Yb [F]) with enrichments in Ba, Rb, Nb, and Th for the Copahue magmas but similar trends for the two groups in Sr. The Nb/Yb values are constant for the southern Southern volcanic zone (SSVZ) and the Caviahue caldera samples, whereas the Riscos Bayos (RB), intracaldera (IC) samples and backarc basalt (BAB) lavas plot slightly above that field. The Copahue lavas plot much higher in the transitional Southern volcanic zone (TSVZ) field and possibly the northern Southern volcanic zone (NSVZ) fields (few data available).
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Geochemistry and isotopic characteristics of the Caviahue-Copahue volcanic complex
D
E
F
Figure 8 (continued).
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A
B
C
Figure 9. Incompatible element diagrams (see Rollinson, 1993, for normalization values) showing the depletion in Ti, P, Nb, and Ta in all rocks, and spikes in U-Th for the most-evolved members (A). The rare earth element (REE) patterns (normalization values after Henderson, 1984) show strong similarities for the whole Copahue-Caviahue volcanic complex (B), including the Riscos Bayos (RB) samples, except for the RB3 rhyolites, which have much steeper light (L) REE patterns (C), similar to the Cerro Bayo rock sample. The two RB3 samples with values below 10 are bulk samples, and the other one is the composite pumice clast sample. BAB—backarc basalt; IC—intracaldera samples.
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TABLE 5. ISOTOPE RATIOS OF CAVIAHUE-COPAHUE VOLCANIC COMPLEX ROCK AND HYDROTHERMAL FLUID SAMPLES AND FOR STANDARD SAMPLE BCR-2 Crater Hot Caviahue Riscos Bayos Copahue Backarc basalts Int'l standard lake spring EB 8 RB1-1 RB1-4 RB2-1 RB3-2 COP 3 AE3D BCR-2 # 1921 87 Sr/86Sr 0.703781 0.703776 0.703783 0.703761 0.703985 0.703954 0.703295 0.704999 143 Nd/144Nd 0.512846 0.512849 0.512854 0.512857 0.512823 0.512795 0.512924 206 Pb/204Pb 18.557 18.554 18.560 18.558 18.594 18.764 207 Pb/204Pb 15.598 15.600 15.602 15.607 15.606 15.623 208 Pb/204Pb 38.435 38.440 38.449 38.461 38.502 38.739 207 Pb/206Pb 0.8394 0.8390 0.840586 0.840750 0.840638 0.840969 0.839283 208 Pb/206Pb 2.0692 2.0702 2.071259 2.071729 2.071597 2.072457 2.070654 Source a a b b b b b b b Note: Sources: a—Varekamp, unpub. data; b—Colvin, 2004. Isotope analyses of Nd, Sr, and Pb were carried out at the Free University in Amsterdam (The Netherlands) using standard protocols of digestion, extraction, and analyses by TIMS and ICP-MS (Sr-Nd measured with a Finnigan MAT262; Pb was measured by MC-ICPMS, Finnigan Neptune). Precision is as follows (2 errors): Sr isotope ratios = 0.000008; Nd isotope ratio – 0.0000075; 2 sigma errors in analyses of standard AGV-1 in Pb isotope ratios: 6/4 = 0.002; 7/4 = 0.001; 8/4 = 0.003. Two Pb isotope analyses were made on 1999 water samples from the Copahue acid hydrothermal fluids (Varekamp et al., 2001) with a single-collector “Finnegan Elements” ICP-MS instrument at URI, Rhode Island, under supervision of Prof. J.G. Schilling. These analyses provide the relative abundances of the 206, 207, and 208 isotopes of Pb, but not 204 (interference with Hg isotopes). All Pb isotope analyses are then recast as 207/206 and 208/206 ratios for display in Figure 14.
Figure 10. The Copahue-Caviahue volcanic complex 87Sr/ 86Sr ratios shown in a regional context. The highest ratios are found in the northern Southern volcanic zone (NSVZ), with much lower values in the transitional Southern volcanic zone (TSVZ) and southern Southern volcanic zone (SSVZ). The Copahue-Caviahue volcanic complex data fit well within this regional pattern, with the RB3 and Copahue samples the highest and backarc basalt (BAB) the lowest in 87Sr/ 86Sr. MORB—mid-oceanridge basalt.
Other isotope diagrams (Colvin, 2004) show similar relations, with the Copahue sample displaced toward the Pacific Ocean detrital sediment field or Nazca plate sediment (NPS) field (e.g., in 143Nd/ 144Nd versus 206Pb/ 204Pb), the Copahue rock plotting among the transitional Southern volcanic zone, and the Caviahue samples plotting in a tight cluster with the southern Southern volcanic zone samples. A similar relationship occurs in the 87Sr/ 86Sr versus 206Pb/ 204Pb, with the Copahue sample displaced toward the Nazca plate sediment field. The He isotope data from two Copahue geothermal pools (Table 6) show similar values of R/Ra of ~7.6, (Ra = air 3He/4He), which is very close to the MORB mantle value (e.g., Hilton et al., 2002; Poreda and Craig, 1989). Studies of Southern volcanic zone hot springs (Hilton et al., 1993) show a range of
values up to 6.6RA, so samples reported here have the highest ratios reported for the region to date. The CO2/ 3He ratio of the fluids is ~9 × 109, which indicates that they are significantly enriched in carbon relative to MORB mantle, indicating a strong slab input (e.g., Varekamp et al., 1992). The carbon isotopic composition of carbon dioxide from bubbling springs is δ13C = –8 ‰, suggesting that carbon of organic origin may be an important contributor. We note, however, that the Copahue geothermal system is rich in organic carbon gases, which could have been generated through thermal processes in the sedimentary basement (e.g., the Neuquén marine sequence at ~2 km depth; Mazzoni and Licitra, 2000). The presence of 129I (iodine) in the Copahue hot-spring fluids is strong evidence for the presence of a sedimentary component rich in organic material 3He/4He
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A
Figure 11. Detail of the Sr-Nd isotope diagram, with the mid-oceanridge basalt (MORB) field, the backarc basalt (BAB) sample, and the Caviahue (open symbols) and Copahue samples (filled circle) displaced toward the Pacific Ocean detrital sediment (PDS) field. Note the tight grouping of the southern Southern volcanic zone (SSVZ) and transitional Southern volcanic zone (TSVZ) rocks on this diagram and the position of the Copahue rock in the transitional Southern volcanic zone field. NSVZ—northern Southern volcanic zone.
B
(Fehn et al., 2002). The acidic Copahue hot-spring fluids are very concentrated magmatic brines (Varekamp et al., 2001, 2004) and are most likely not impacted by the presence of shallowlevel sedimentary rocks. DISCUSSION The Copahue-Caviahue volcanic complex rocks consist of the older Caviahue series of lavas (caldera walls) that are basaltic andesites to dacites with intercalated more-silicic pyroclastic deposits, the younger Riscos Bayos and intracaldera rocks, and the young Cerro Bayo silicic dome. The Copahue series, made largely of basaltic andesites, postdates this sequence and is farther to the west, slightly closer to, but still east of the modern volcanic front. The backarc basalts are relatively young and lie outside the main Caviahue volcanic complex to the east in a broad band of cinder cones and basaltic lava fields (Fig. 2). Based on data presented here, some conclusions can be drawn on the relations of these units. Contrary to earlier work, which suggested that the Riscos Bayos units consisted of a single rock type, the Riscos Bayos ignimbrites consist of at least three major pumice-rich flows, with an underlying mud-rich flow. The RB1 flow has white pumices, black obsidian and small dark pumice fragments with the same composition as some pumice clasts in RB2. The RB1 and RB2 flows were deposited during a continuous volcanic event, where RB2 was deposited later than RB1. The RB3 unit is distinct in its mineralogy (quartz + biotite), is the most-evolved rock in the region (rhyo-
Figure 12. (A) Overview and (B) detail of the 207Pb/ 204Pb versus 206Pb/ 204Pb, showing the displacement of the Copahue-Caviahue volcanic complex (CCVC) samples away from the mid-ocean-ridge basalt (MORB) and Nazca plate basalt (NPB) fields toward the field of Pacific detrital sediments (PDS). Both groups plot inside the sediment field, and the Copahue sample has the most radiogenic isotope composition. NHRL—Northern Hemisphere reference line.
lite), is strongly enriched in several trace elements, especially U and Th, and is geochemically similar to the younger Cerro Bayo dome. The REE patterns of both RB3 and the Cerro Bayo dome rock show a flat HREE and very steep LREE trend. The RB3 unit has the most radiogenic Pb and Sr isotopic ratios, and the lowest Nd isotopic ratio of the Caviahue rock samples. Taken together, the RB3 magma carries an upper-crustal imprint, possibly acquired under an AFC scenario with assimilation of U, Th, and their radiogenic daughters.. The Riscos Bayos sequence may represent an inverted zoned magma chamber, where the silicic top, with a strong crustal imprint, erupted first, followed by the less-silicic RB1 and RB2 flows.
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B
A
Figure 13. (A) Overview and (B) detail of the 208Pb/ 204Pb versus 207Pb/ 204Pb diagram. The Caviahue samples plot just outside the Pacific Ocean detrital sediment field in a linear array, with RB3 as the most contaminated sample. The Copahue sample plots just inside the sediment field and off the Caviahue sample array. CCVC—Copahue-Caviahue volcanic complex; MORB—mid-ocean-ridge basalt; RB—Riscos Bayos.
Figure 14. Detail of the 208Pb/ 206Pb versus 207Pb/ 206Pb diagram. The x-axis represents the ratio of the two U daughters, whereas the y-axis represents the ratio of the Th daughter and one of the U daughters. The two samples from hot-spring fluids at Copahue plot close to the Copahue rock sample, confirming the anomalous composition of Copahue relative to Caviahue samples. The Caviahue samples may contain an “older” sediment component than the Copahue rocks. MORB— mid-ocean-ridge basalt.
The isotopic differences between the mafic and more evolved Caviahue units are very small, suggesting that the amount of assimilation during magmatic evolution was small and/or that the crustal component was not very different in isotopic composition from the magmas themselves. The Copahue rocks are enriched in incompatible elements, both large ion lithophile elements (LILEs) and high field strength elements
(HFSEs), relative to the older Caviahue rocks, but depleted in Ca and Fe. Their isotopic signature differs from Caviahue samples with significantly higher 206Pb/ 204Pb and 208Pb/ 204Pb values, and, for such mafic rocks, very high Sr isotope ratios (almost the same ratios as those found in the silicic RB3 pumices) and lower Nd isotope ratios. In terms of many major elements, trace elements, and isotopic ratios, the Copahue rocks have strong affinities with
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TABLE 6. HELIUM AND CARBON ISOTOPE DATA FROM GEOTHERMAL FLUIDS AT COPAHUE VOLCANO 13 Sample RM/RA † X‡ RC/RA§ CO2/3He C (×109) Baño 7 7.6 1228 7.6 ± 0.1 9.8 ± 0.2 –8.28 + 0.01 Las Máquinas 7.5 729 7.5 ± 0.1 9.0 ± 0.1 –8.32 + 0.01 Note: Carbon and helium isotope data were determined on gases collected from bubbling, low-temperature ponds in the geothermal area near the village of Copahue (Las Máquinas pool, January 2003; spa pool #7 in Copahue village, January 2003). Inverted funnels were placed underwater, and the gases were let into evacuated lead-glass bottles. The He and C isotope analyses were carried out according to protocols described by Shaw et al. (2003). † RM/RA = measured 3He/4He (RM) in sample relative to air 3He/4He (RA). ‡ X = [(He/Ne)sample/(He/Ne)air] 1.209 (i.e. the air-normalized He/Ne ratio multiplied by Ne/He = 1.209 – the ratio of the Bunsen coefficients at 17 °C; Hilton, 1996). § RC/RA = air-corrected 3He/4He (RC) in sample relative to air (RA)—see Hilton (1996) for details of the correction procedure.
the transitional and northern Southern volcanic zone volcanoes, whereas the Caviahue rocks show more similarities with the rocks from the southern Southern volcanic zone. Magma Genesis in the Copahue-Caviahue Volcanic Complex Area The chemical and isotopic data from the Copahue-Caviahue volcanic complex rocks indicate that all the rocks have undergone some form of enrichment in incompatible elements compared to pure mantle melts, including the backarc basalts. The normalized incompatible element concentrations of the CopahueCaviahue volcanic complex rocks show enrichment with respect to normal (N)-MORB (Fig. 15A). The RB3 pattern resembles that of the upper crust with additional enrichment in U and Th. All Copahue-Caviahue volcanic complex rocks show enrichments in LREE relative to the Nazca plate basalts
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B
Figure 15. Source diagrams for the incompatible elements. Note the strong U-Th enrichment in RB3 relative to upper continental crust (Rollinson, 1993), and the increasing enrichment in very incompatible elements in the more evolved rocks (A). The rare earth element (REE) diagram (B) shows the anomalous composition of the RB3 samples and the overall enrichment in light (L) REEs relative to the Nazca plate basalts. RB— Riscos Bayos ignimbrite units; BAB—backarc basalt; IC—intracaldera samples; N-MORB—normal mid-ocean-ridge basalt
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Geochemistry and isotopic characteristics of the Caviahue-Copahue volcanic complex
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B
Figure 16. Potential evidence for crustal assimilation during crustal residence shown by the K/Rb versus Rb concentrations in Copahue-Caviahue volcanic complex (CVCC) and Southern volcanic zone rocks (A). The field for fractional crystallization with the backarc basalt (BAB) as a parental magma is after Davidson et al. (1987). Increases in Sr isotope ratios with decreasing K/Rb are common in the northern Southern volcanic zone rocks (NSVZ) (B), whereas only a weak relationship exists for southern Southern volcanic zone (SSVZ) and transitional Southern volcanic zone (TSVZ) rocks. The Copahue-Caviahue volcanic complex samples show a slight increase in Sr isotope ratios with decreasing K/Rb. PDS— Pacific Ocean detrital sediment; IC—intracaldera, RB—Riscos Bayos; UC—upper crust.
(Fig. 15B), with very similar REE patterns for all Caviahue and Copahue rocks, with the exception of RB3 and Cerro Bayo dome, which have a much steeper LREE pattern. The source of the material rich in incompatible elements could be subducted sediment and/or slab components with “arc enrichment” as a result of volatile fluxing of the subducted complex or melting of the subducted complex. In addition, slices of eroded material from the overriding plate may have been dragged down (e.g., Stern, 1991), and crustal assimilation during crustal residence may have occurred. The relation between Rb concentrations and K/Rb has been used to estimate the extent of crustal contamination in evolving Southern volcanic zone magmas (e.g., Davidson et al., 1987), and the Copahue-Caviahue volcanic complex rocks plot within the trend observed for other Southern volcanic zone rocks (Fig. 16A). The field for closedsystem fractional crystallization using the backarc basalt as a hypothetical parent is indicated, and from this graph as well as the isotopic data, it is obvious that the Copahue-Caviahue volcanic complex suite cannot have been derived from fractional crystallization alone from a backarc basalt–type parental magma. The K/Rb versus 87Sr/ 86Sr data show that in the northern Southern volcanic zone, the Sr isotope ratios increase with decreasing K/Rb, whereas in the southern and transitional Southern volcanic zone, this is not the case (Fig. 16B). The Copahue-Caviahue volcanic complex samples show a slight increase in Sr isotope ratios with decreasing K/Rb (Fig. 16B), and we conclude that the crustal rocks that were assimilated during crustal residence have only a small isotopic difference from the parental Copahue-Caviahue volcanic complex magmas. Detailed AFC modeling is needed to derive a possible range of compositions of the crustal assimilant and/or the composition of the parental magmas. If the parental magmas are closer in chem-
ical and isotopic composition to the mafic caldera wall lavas (e.g., such as EB8), there is little reason to invoke massive amounts of crustal assimilation (except for RB3), because most of the K/Rb data can then be explained by largely closed-system fractional crystallization. Another mechanism must then have caused the relatively radiogenic isotopic composition of these parental magmas, e.g., a larger component extracted from the subducted complex. With both uncertainty in the composition of the parental magmas and the crustal assimilant, quantitative AFC modeling will not provide much new insight at this point. The chemical differences between the Copahue and Caviahue rocks are best summarized as follows: Copahue rocks are enriched in both LILE and HFSE relative to the Caviahue series and are isotopically characterized by higher Sr and Pb isotope ratios with anomalously high 208Pb/ 204Pb, 206Pb/ 204Pb, and lower Nd isotope ratios. They are also among the more mafic rocks in the southern Southern volcanic zone, and the measured He isotope ratios are close to mantle values. If the trace-element and isotopic trends within the Caviahue suite are best explained as resulting from open-system evolution in the continental crust, then the Copahue rocks are unlikely to represent a more advanced degree of that same process. The Nb/Yb values are constant in the whole Caviahue series (Fig. 8F), whereas the Copahue samples plot above that array at similar MgO contents. The AFC processes that were active during development of the Caviahue series did not lead to an increase of that element ratio, and such AFC processes are thus not a likely cause for the higher Nb/Yb in the Copahue rocks. If enhanced crustal assimilation is the reason for the chemical and isotopic characteristics of the Copahue series, most likely that section of the crust had a different chemical and isotopic composition (e.g., lower Nd isotope ratios) than the crust that influenced the Caviahue volcanic series.
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Figure 17. Trace-element ratio diagram showing mid-ocean-ridge basalt (MORB), upper continental crust (UCC) (after Rollinson, 1993), Pacific Ocean detrital sediment (PDS), and the Copahue-Caviahue volcanic complex (CVCC) data together with the Southern volcanic zone data set. The trend for arc enrichment as a result of fluid fluxing of subducted sediments is indicated along the vertical axis, whereas the degree of melting of sediments or continental crust is shown along the horizontal axis. The southern Southern volcanic zone (SSVZ) samples plot at the upper left and represent arc enrichment by fluid fluxing. The Caviahue samples plot at the bottom of that field, except for the RB3 samples, which plot close to continental crust and Pacific Ocean detrital sediment. The transitional Southern volcanic zone (TSVZ) and northern Southern volcanic zone (NSVZ) samples show evidence for crustal melting. The Copahue samples plot below all other Southern volcanic zone rocks as a result of very low Ba/La values. These magmas may have resulted from melting of a fractionated sediment residue in the subducted complex that lost volatiles and Ba during an earlier fluxing stage (e.g., below the modern Southern volcanic zone arc). RB— Riscos Bayos ignimbrite units; BAB—backarc basalt; IC—intracaldera samples.
An alternative possibility is the difference in mode of acquisition of incompatible elements from the sources for the Copahue series: sediment melting versus fluid fluxing of the subducted complex. The former would explain the enrichment of HFSE in the Copahue magmas. The He isotopic values suggest mainly a mantle source for the He, in apparent contrast with the evidence from the other isotope data for the presence of a crustal component in the Copahue magmas. Loss of radiogenic He “higher up” in the subduction zone, e.g., below the main arc to the west of Copahue, may have occurred. An earlier degassing episode of the subducted complex would also agree well with the relatively water-poor character of the Copahue magmas (melt inclusions in olivine with 1%–2% H2O; Goss, 2001) and lack of large-scale explosive activity. A similar process was discussed by Hilton et al. (1992) and Hoogewerff et al. (1997) in Indonesia, who suggested He loss from the subducted complex “higher up” in the subduction zone, resulting in a smaller contribution of radiogenic, sediment-derived He in the more distal sections of that subduction zone.
The La/Yb versus Ba/La diagram (Cameron et al., 2003; Fig. 17) may show the differences that occur as a result of slab fluxing by fluids (increasing Ba/La) versus melting processes. The La/Yb value increases with addition of crustal material to mantle melts, and it increases with decreasing percentage of melting of source rocks, be it mantle or crustal material. The southern Southern volcanic zone rocks show the strongest evidence for “arc enrichment” by slab fluxing of the Southern volcanic zone, and weaker signals are observed northward. Instead, in the transitional and northern Southern volcanic zone, the incorporation of crustal material seems to become more prevalent (larger La/Yb), as shown by the work of Hildreth and Moorbath (1988). The Sr isotope versus latitude diagram (Fig. 10) suggests the same, and this process of enhanced contributions of continental crust may be related to the increasing age and thickness of the crust to the north (Davidson et al., 1987; Hildreth and Moorbath, 1988). The Copahue rocks and backarc basalt show little evidence of this “arc-enrichment” process by fluid fluxing and plot at lower Sr isotope values than most other rocks of the Southern volcanic zone, but at higher La/Yb values. The low Ba/La values may be explained as a result of melting of a sediment residue that had already lost Ba during an earlier fluid fluxing phase, e.g., below the main modern Southern volcanic zone arc. The Caviahue samples plot within the southern Southern volcanic zone array, which seems to be dominated by fluid fluxing of the subducted complex and low La/Yb values, typical for higher degrees of melting associated with flux melting (Cameron et al., 2003) and low contributions of crustal material with high La/Yb. The RB3 samples plot near bulk crustal material (upper continental crust and Pacific detrital sediment), suggesting addition of crustal material to the evolving magmas during residence in the continental crust. The higher La/Yb values in the Copahue and backarc basalt samples compared to the Caviahue rocks point to addition of crustal material and possibly also to lower degrees of partial melting (Bermúdez et al., 2002). The Copahue magmas may have formed as a result of low degrees of partial melting of a mantle wedge contaminated with melts of subducted sediment. This subducted sediment had already lost a fraction of its volatiles and water-soluble elements below the modern Southern volcanic zone active volcanic front to the west, as indicated by high He isotope ratios and low Ba/La values. The reduced amount of fluid fluxing led to smaller degrees of partial melting in this drier mantle environment, leading to the higher La/Yb. The magmatism of Copahue and the backarc basalt cinder cones in the Loncopué graben may be related to decompressional melting (e.g., Cameron et al., 2003) related to the modern phase of extensional tectonics (Bermúdez et al., 2002). The He-CO2 relationships at Copahue suggest the presence of slab-derived fluids in the geothermal gases, and adopting the approach of Sano and Marty (1995), it can be calculated that ~91% of the carbon originates from the subducted complex, with the remainder (9%) coming form the mantle wedge. The impact of shallow carbon-bearing sediments below the volcano cannot be discounted, however.
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Geochemistry and isotopic characteristics of the Caviahue-Copahue volcanic complex In conclusion, many of the regional isotope and traceelement trends in the Southern volcanic zone may be the result of changes in the upper plate, as argued by Hildreth and Moorbath (1988). To explain the differences in the Copahue-Caviahue volcanic complex, we either have to invoke crustal contaminants with variable chemical and isotopic compositions or variations in the extraction process from the subducted sediments. The latter would be associated with changes in mantle melting processes, as well as a reflection of earlier depletion in fluid mobile elements, when the subducted complex passed below the main modern active Southern volcanic zone arc. The difference between fluid fluxing and sediment melting may also impact the isotopic characteristics of the resulting melts, given the relatively low solubility of Th-bearing compounds in fluids. Magmas with sediment melt would have higher Th concentrations and possibly enhanced 208Pb/ 204Pb, as found in the Copahue series. CONCLUSIONS Volcanism in the Copahue-Caviahue volcanic complex over the last 4 m.y. has evolved through two separate broad stages: (1) establishment of the Caviahue volcano, with subsequent caldera formation, and (2) the formation of Copahue volcano over the last 0.8 m.y. The backarc cinder cones in the Loncopué graben formed relatively recently and are isotopically and geochemically among the least-evolved arc-related rocks in the Southern volcanic zone. New findings regarding the local volcanological evolution are: (1) The caldera walls consist of lavas and ignimbrites that do not show a gradual evolution toward a climactic phase. In fact, the most mafic lavas are found in the top one-third of the sequence of the north and east caldera walls. (2) The Riscos Bayos ignimbrites consist of at least four stratigraphic units, with three major pumice-rich flows of distinct chemical composition and mineralogy. Their volume is insufficient to explain the formation of the Caviahue caldera. The oldest unit (RB3) shows chemical and isotopic evidence for crustal assimilation processes. (3) The intracaldera silicic flow complex on the Lake Caviahue peninsula has mineralogy and bulk chemical composition that are distinct from those of the Riscos Bayos, and, as such, it is unlikely that it is the intracaldera facies of the Riscos Bayos flows (Mazzoni and Licitra, 2000). (4) The Cerro Bayo dome rock has a similar bulk composition to the RB3 flow, but a different mineralogical constitution and is much younger than RB3 as well. (5) All volcanic rocks from Copahue-Caviahue volcanic complex carry evidence of a continental source component in their isotopic signatures. The chemical and isotopic fingerprint of the Copahue rocks is distinct from that of the Caviahue series; notably, the Copahue rocks are more enriched in LILEs and HFSEs than the Caviahue rocks. (6) The Copahue rocks have similarities with the transitional and northern Southern volcanic zone rocks, whereas the Caviahue rocks resemble those from the southern Southern volcanic zone.
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The establishment of Copahue volcano was associated with a change in magma source components. This could be explained ad hoc as a result of AFC processes with a different section of continental crust than that involved with the Caviahue magmatic series. Alternatively, we propose that the mechanism of element extraction from the subducted complex may have changed from fluid fluxing to largely partial melting, as indicated by the trace-element data. The He isotope and traceelement data from Copahue also suggest that the subducted sediment may have become fractionated as a result of earlier fluid extraction processes below the active volcanoes to the west. The change from Caviahue to Copahue magmatism may then be related to a broad change from a fluid fluxing regime several million years ago to a sediment melting regime for the last million years. A tentative explanation for this change is the following: the Nazca plate has a number of old sutures that juxtapose older and younger oceanic crust (e.g., the Mocha fracture zone). Subduction of older and, therefore, colder ocean-floor sections below the Copahue-Caviahue volcanic complex led to the Caviahue volcanism, with fluid fluxing as the primary “arc enrichment” mechanism. Around 1 Ma, a suture may have passed below the Copahue-Caviahue volcanic complex, and a warmer section of the Nazca plate was subducted. As a result, volatile release occurred earlier in the subduction process (below the main modern Southern volcanic zone arc west of Caviahue, Fig. 1), and a fractionated residue of the subducted sediment appeared below the Copahue-Caviahue volcanic complex zone, where sediment melting started to occur in the zones of magma generation (Copahue magmatism). We cannot exclude that sediment on the oceanic crust with a different age may have had a different chemical and isotopic composition, which could explain the distinct isotopic composition of the Copahue series. The backarc basalts received the most fractionated sediment contribution, but the total contribution was small, providing only minor isotopic shifts away from the MORB mantle composition. The nature of mantle melting may have switched from flux melting to decompressional melting with extensional tectonics in the region. The more-evolved Copahue-Caviahue volcanic complex rocks are very similar to average upper-crustal rock: traceelement patterns normalized to the upper crust or average North American shales have element ratios close to 1 both for REEs (this study; Gammons et al., 2005) and for other incompatible trace elements, with U and Th somewhat enriched above normal crustal levels. ACKNOWLEDGMENTS Funding for this research was provided by grants from the National Science Foundation (NSF) (INT-9704200 and INT9813912), the National Geographic Society (grant 7409–03), and support from the Smith Fund at Wesleyan University. Daniel Delpino and Adriana Bermúdez have been our long-time collaborators in our studies of Copahue volcano, and we are
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grateful for their contributions in the field and help with logistics over the years. Much of the work presented here is based on senior thesis research of Adam Goss, Maarten deMoor, Matt Merrill, and Anna Colvin, when they were undergraduates at Wesleyan University. Reviews by Robert Kay and Jon Davidson were very helpful, but remaining unsubstantiated speculations are the full responsibility of the authors. REFERENCES CITED Bachmannn, O., Dungan, M.A., and Lipman, P.W., 2000, Voluminous lava-like precursor to a major ash-flow tuff: Low-column pyroclastic eruption of the Pagosa Peak Dacite, San Juan volcanic field, Colorado: Journal of Volcanology and Geothermal Research, v. 98, p. 153–171, doi: 10.1016/ S0377-0273(99)00185-7. Barazangi, M., and Isacks, B.L., 1976, Spatial distribution of earthquakes and subduction of the Nazca plate beneath South America: Geology, v. 4, p. 686–692, doi: 10.1130/0091-7613(1976)4<686:SDOEAS>2.0.CO;2. Beck, S.L., Zandt, G., Myers, S.C., Wallace, T.C., Silver, P.G., and Drake, L., 1996, Crustal-thickness variation in the central Andes: Geology, v. 24, no. 5, p. 407–410, doi: 10.1130/0091-7613(1996)024<0407:CTVITC>2.3.CO;2. Bermúdez, A.M., and Delpino, D., 1999, Erupciones subglaciales y en contacto con hielo en la región volcánica de Copahue, Neuquén, Argentina, in XIV Congreso Geológico Argentino (Salta): Buenos Aires, Actas II, p. 250–253. Bermúdez, A.M. and Delpino, D.H., 2002. Las erupciones del volcán Copahue del año 2000. Impacto social y del medio económico. Provincia del Neuquén, Argentina. 15° Congreso Geológico Argentino (El Calafate), Actas, 3: 365-370, Buenos Aires. Bermudez, A., Delpino, D., Frey, F., and Saal, A., 1993, Los basaltos de retroarco extraandinos, in Ramos, V.A., ed., Geología y Recursos Naturales de Mendoza, Relatorio, XII° Congreso Geológico Argentino (Mendoza): Buenos Aires, p. 173–195, Bermúdez, A., Delpino, D., and López-Escobar, L., 2002, Caracterización geoquímica de lavas y piroclastos holocenos del volcán Copahue, incluyendo los originados en la erupción del año 2000. Comparación con otros volcanes de la Zona Volcánica Sur de los Andes, in Congreso Geológico Argentino, No. 15: El Calafate, Actas, v. 1, p. 377–382. Bevis, M., and Isacks, B.L., 1984, Hypocentral trend surface analysis; probing the geometry of Benioff zones: Journal of Geophysical Research, v. 89, no. 7, p. 6153–6170. Cahill, T., and Isacks, B.L., 1992, Seismicity and shape of the subducted Nazca plate: Journal of Geophysical Research, v. 97, B12, p. 17,503–17,529. Cameron, B.I., Walker, J.A., Carr, M.J., Patino, L.C., Matias, O., and Feigenson, M.D., 2003, Flux versus decompression melting at stratovolcanoes in southeastern Guatemala: Journal of Volcanology and Geothermal Research, v. 119, p. 21–50, doi: 10.1016/S0377-0273(02)00304-9. Cembrano, J., Hervé, F., Lavenu, A., 1996, The Liquiñe-Ofqui fault zone; a long-lived intra-arc fault system in southern Chile: Geodynamics of the Andes: Tectonophysics, v. 259, no. 1–3, p. 55–66. Colvin, A.S., 2004, Trace element and isotope geochemistry of the CaviahueCopahue volcanic complex [Undergraduate thesis]: Middletown, Connecticut, Wesleyan University, p. 207. Davidson, J.P., Dungan, M.A., Ferguson, K.M., and Coluccii, M.T., 1987, Crust magma interactions and the evolution of arc magmas: The San Pedro– Pellado volcanic complex, southern Chilean Andes: Geology, v. 15, p. 443–446, doi: 10.1130/0091-7613(1987)15<443:CIATEO>2.0.CO;2. Davidson, J.P., Ferguson, K.M., Colucci, M.T., and Dungan, M.A., 1988, The origin and evolution of magmas from the San Pedro–Pellado volcanic complex, S. Chile: Multi-component sources and open system evolution: Contributions to Mineralogy and Petrology, v. 100, no. 4, p. 429–445, doi: 10.1007/BF00371373.
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