Abstract Many areas of both California and the Great Basin contain diatom-bearing sediments which range in age from Late Cretaceous to Recent. The diatom assemblages in these rocks contain both short-ranging species that are useful for stratigraphic correlation and others, still represented in living assemblages elsewhere, that are useful for paleoecologic interpretations. Although thousands of square miles of diatom-bearing sediments have been mapped by petroleum geologists and have been penetrated and cored during drilling operations, the diatoms have been neglected as a stratigraphie and paleoecologic tool. Some of the reasons most often presented for this lack of attention are claims that Foraminifera are easier to work with, are better known, and that laboratory manipulation of them is easier and more suited to assembly-line methods. This is true in part only, as assembly-line methods have been developed for handling large numbers of samples of diatomaceous sediments. Furthermore, diatoms are often found in sediments that are completely barren of Foraminifera or other fossils. The present interest in palynology by the oil companies indicates that new techniques are no longer received with disfavor. As far as the necessary laboratory preparation and study are concerned, diatom samples can be prepared and significant species identified at least as expeditiously as samples containing pollen and spores. Distinctive diatom assemblages are known from the Moreno Shale of Late Cretaceous and Paleocene (?) age and from many sedimentary formations in Eocene, Oligocene, Miocene, Pliocene, and Pleistocene rocks in California from the San Francisco Bay area southward. These assemblages from rocks of Cretaceous through Miocene age are virtually all marine. Pliocene rocks in different localities contain either marine or nonmarine diatom assemblages. Pleistocene asemblages are dominantly nonmarine. Extensive areas of Miocene, Pliocene, and Pleistocene sediments in Nevada and other parts of the Great Basin also contain distinctive nonmarine diatom assemblages. In that region, diatoms are often the only fossils present. Here also, the diatoms can provide much needed paleoecological information, as the Cenozoic lake basins varied greatly in depth, temperature, salinity, pH, and other factors of paleoecological importance.

Abstract Magmatism in the Chilean Andes has taken place since about 300 Ma as a consequence of protracted subduction, although with significant spatial and temporal variations due to changes in ocean-floor geodynamics controlling distinct large-scale magmatic events. Early subduction along the Chilean segment of the Gondwana active margin took place during Late Palaeozoic times and generated typical arc magmatism and a subduction complex in the forearc environment. This tectonomagmatic regime was interrupted by mid-Permian contractional tectonics (collisional?) giving rise to a thickening of the crust that allowed deep crustal melt generation. Following this the entire Mesozoic history of the area became dominated by subduction-related extensional tectonics with mostly bimodal magmatism reflecting the involvement, to different degrees, of both crust and mantle as magma sources. Mesozoic volcanism and plutonism appear to have been independent of each other. Subsequent Cenozoic magmatism records changing geodynamic conditions from Palaeogene–early Neogene extension to late Neogene compression. The Neogene magmatic episodes are interpreted as an indirect consequence of oceanic ridge subduction: the Juan Fernández Ridge along the north-central Chilean margin, and the Chile Ridge along the southernmost Chilean border. Modern volcanism is also influenced by these ridge subductions, either by generating gaps in the Quaternary volcanic chain, or adakitic volcanism derived from slab melting. Despite the essentially tectonic control outlined above, this chapter is subdivided geographically into four Andean segments, each of which exhibits distinct magmatic features. These segments are: 18–28°S, 28–38°S, 40–47°S and 47–55°S. The exception to this approach is the section on Quaternary volcanism, which

Abstract The Canarian archipelago comprises seven main volcanic islands and several islets that form a chain extending for c . 500 km across the eastern Atlantic, with its eastern edge only 100 km from the NW African coast (Fig. 18.1 ). The islands have had a very long volcanic history, with formations over 20 million years old cropping out in the eastern Canaries. Thus all stages of the volcanic evolution of oceanic islands, including the submarine stage as well as the deep structure of the volcanoes, can be readily observed. Rainfall and vegetation cover are relatively low, with the exception of the island of La Palma, favouring both geological observation and rock preservation. Furthermore, the absence of surface water has promoted groundwater mining by means of up to 3000 km of subhorizontal tunnels (locally known as ‘galerías’). These galerías are especially numerous in Tenerife, La Palma and El Hierro, and allow the direct observation and sampling of the deep structure of the island volcanoes without requiring expensive and indirect geophysical methods ( Carracedo 1994 , 1996a , b ). Since the early work of famous naturalists such as Leopold von Buch, Charles Lyell, and Georg Hartung, the Canaries have been viewed as a ‘special’ volcanic island group and their origin has been closely related to African continental tectonics ( Fúster et al . 1968a , b , c , d ; McFarlane & Ridley 1969 ; Anguita & Hernán 1975 ; Grunau et al . 1975 ). However, a

Abstract Large slope failures in fractured rocks are often controlled by the combination of pre-existing tectonic fracturing and brittle failure propagation in the intact rock mass during the pre-failure phase. This study focuses on the influence of fold-related fractures and of post-folding fractures on slope instabilities with emphasis on Turtle Mountain, located in SW Alberta (Canada). The structural features of Turtle Mountain, especially to the south of the 1903 Frank Slide, were investigated using a high-resolution digital elevation model combined with a detailed field survey. These investigations allowed the identification of six main discontinuity sets influencing the slope instability and surface morphology. According to the different deformation phases affecting the area, the potential origin of the detected fractures was assessed. Three discontinuity sets are correlated with the folding phase and the others with post-folding movements. In order to characterize the rock mass quality in the different portions of the Turtle Mountain anticline, the geological strength index (GSI) has been estimated. The GSI results show a decrease in rock mass quality approaching the fold hinge area due to higher fracture persistence and higher weathering. These observations allow us to propose a model for the potential failure mechanisms related to fold structures.

Abstract Chile is, geographically, an unusual and in many ways astonishing country ( Fig. 1.1 ). It stretches north–south along the South American mainland for over 4000 km, from 18°S, where the Altiplano is shared with Peru, Bolivia and Argentina, to 56°S at Tierra del Fuego and the islands of Cape Horn, the next stop being Antarctica. Its western margin everywhere is the Pacific Ocean, and its eastern boundary is the summit of the Andes mountains, so that in a width of rarely more than 200 km, the topography rises from sea level to a maximum of almost seven thousand metres. Climatic variations reflect this extraordinary topography. The north is characterized by the Atacama Desert, considered to be the driest place on Earth. The south is in the temperate rainforest zone, with vegetation that struggles against the prevailing westerly gales. In this southern sector the land is moulded by recent glaciations that carved the coastal areas into fiords and archipelagos consisting of thousands of islands; the length of the Chilean coastline including these islands must exceed that of many other countries that have a larger surface area. It is the extreme variety represented by these factors that have led to Chile becoming such an attractive tourist destination, despite the isolation and comparative difficulty of access of many of its geomorphological treasures. Figure 1.1 shows the distribution of the main tectonic and geomorphological features of Chile. The northern and central parts of the country can be reasonably divided into three north– south

Abstract The present-day Andes have formed in response to subduction-related processes operating continuously along the western margin of South America since the Jurassic period. When these processes started, the continental margin was mainly formed of metamorphic complexes and associated magmatic rocks which evolved during Proterozoic (?), Palaeozoic and Triassic times, and which now constitute the basement to the Mesozoic and Cenozoic Andean sequences. These older units are commonly referred to in the Chilean geological literature as the ‘basement’ or the ‘crystalline basement’. The basement rocks crop out discontinuously ( Fig. 2.1 ) in northern Chile, both in the coastal areas and in the main cordillera. In contrast, from latitude 34°S southwards, they form an almost continuous belt within the Coastal Cordillera extending to the Strait of Magellan. In addition, sparse outcrops occur both in the main Andean cordillera as well as further east in the Aysen and Magallanes regions. In the first maps and syntheses of the geology of Chile (e.g. Ruiz 1965 ) these rocks were generally considered to be of Precambrian age, forming a western continuation of the Brasilian craton. Later work has demonstrated that rocks first described as metamorphic basement units show a wide range of metamorphic grades and ages extending from possible Late Proterozoic through Palaeozoic and even, in some cases, to Jurassic–Cretaceous. With regard to previous works that have attempted to synthesize data on Chilean basement geology, the reader is referred to those by González-Bonorino ( 1970 , 1971 ), González-Bonorino & Aguirre (1970) , Aguirre et al .

Abstract This chapter describes the metallic ore deposits of Chile, their mineralized host rocks and the processes involved in ore formation, and provides a brief overview of the mining history of this Andean copper-rich country. The ore deposits are ordered according to their respective economic importance. Thus, after mining history and a general introduction, Chilean porphyry copper–molybdenum deposits are described first, with subsequent sections dealing with epithermal precious metals, iron oxide copper–gold and iron oxide–apatite deposits, stratabound copper–(silver) ores, precious metal veins, sedimentary-hosted gold and porphyry gold deposits, skarn rock ores and, finally, an overview of metallogenic evolution. About 40% of the known copper resources of the world occur in Chile, with the native populations using the red metal at least since 500 BC. Bracelets, earrings and weapons that have been found in archaeological sites in northern Chile were made of either native copper or copper-rich minerals that were melted in small quantities and subsequently hammered. Copper production during Spanish colonial times (1541–1810) amounted to some 80 000–85 000 tons, with high-grade oxidized copper minerals being exploited and melted with charcoal. Despite this mining activity, however, Spaniards regarded copper as ‘plebeian metal’ because of its relatively low value, and it was used mostly as ballast for ships returning to Spain, rather than for technological or industrial purposes. The colonial Spaniards were much more interested in gold and silver, and mining activities were consequently mostly orientated towards precious metals. Prior to Spanish conquest the Incas dominated northern Chile and had already exploited

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Abstract The Devonian was one of the first Palaeozoic periods to be intensively studied in Spain. A few years after the formal definition of the Devonian by A. Sedgwick and R. I. Murchison in Devon, the French naturalists E. de Verneuil and A. d’Archiac (1845) noticed the occurrence of Devonian shelly fossil faunas in Asturias (north Spain). Later on, Prado & Verneuil (1850) enlarged the known Devonian outcrop area to the neighbouring province of Leon, and Prado (1856) extended this to Palencia province. Verneuil & Collomb (1853) , Verneuil & Lorière (1854) and Verneuil & Lartet (1863) demonstrated Devonian rocks in the Iberian Ranges, and both Almera (1891c) and Barrois (1892) were pioneers in the study of Devonian rocks in the Catalonian Coastal Ranges. In southern Spain the seminal work on the system belongs to E. de Verneuil and J. Barrande ( Prado et al. 1855 ), and in the Balearic Islands Hermite (1879) discovered the Devonian succession of Minorca. The history of Devonian research in other Spanish areas is in general much more recent, and was mainly developed in the twentieth century ( Julivert et al. 1983 ). Devonian rocks everywhere in Spain were deposited in marine conditions, although in varied settings ranging from supratidal to subtidal environments. The thickest and most complete Devonian succession in Spain is found in the Cantabrian and WestAsturo-Leonian zones and in the Basque Pyrenees (a–f and w,